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LATE HOLOCENE CLIMATE FLUCTUATIONS AT CASCADE LAKE,
NORTHEASTERN AHKLUN MOUNTAINS, SOUTHWESTERN ALASKA
By Kasey Kathan
A Thesis
Submitted in Partial Fulfillment
of the Requirement for the Degree of
Master of Science
in Geology
Northern Arizona University
December 2006
Approved:
________________________________
Darrell S. Kaufman, Ph.D., Chair
_________________________________
R. Scott Anderson, Ph.D.
_________________________________
James Sample, Ph.D.
_________________________________
Al Werner, Ph.D.
ii
ABSTRACT
LATE HOLOCENE CLIMATE FLUCTUATIONS AT CASCADE LAKE,
NORTHEASTERN AHKLUN MOUNTAINS, SOUTHWESTERN ALASKA
KASEY KATHAN
This study examines the relationships between glacier fluctuations, multiple
proxies from lake sediments and late Holocene climate fluctuations in southwestern
Alaska. Cascade Lake (3.1 km2; 60.1°N; 159.4°W; 330 m asl) is glacier-fed with 12% of
its drainage basin occupied by modern (1970) glaciers. The primary sediment contributor
is Cascade Glacier located 5.4 km up-valley. During the Little Ice Age, cirque and valley
glaciers of the Cascade Lake area experienced an average equilibrium-line altitude
lowering of 22 ± 12 m (n = 8) relative to 1957. Little Ice Age moraines stabilized ~1860
AD.
Three sediment cores were recovered from the lake and the lake was instrumented
with water-temperature loggers and sediment traps from 2004-2006. 137Cs, 210Pb profiles,
the Aniakchak tephra (3500 cal yr BP) and the regional tephra (Tephra B; 6100 cal yr
BP) provide limited age control and indicate an average late Holocene sedimentation rate
of ~0.07 cm/yr. The cores are composed of rhythmically laminated mud, with an average
lamination thickness of 0.4 cm; these laminations do not represent annual sedimentation
(e.g. varves).
Biogenic silica was analyzed at an average spacing of 0.2 cm (~3 yr) over the last
150 yr and is well correlated with spring temperature (r = 0.70, p = 9.16 x 10-5),
precipitation (Nuyakuk discharge; r = 0.73, p = 2.2 x 10-6) and the Aleutian Low Pacific
iii
Index (r = 0.68, p = 1.37 x 10-7). These relationships suggest a link to nutrient input to
the lake. High BSi flux can generally been related to mild conditions and low BSi flux to
cool conditions. This interpretation and other proxy data indicate mild climates at ~7000
cal yr BP (Holocene thermal maximum) and 1600-1400 cal yr BP (Medieval Warm
Period?). Cool conditions persisted between ~3000-2500 cal yr BP (onset of
Neoglaciation) and 1400 cal yr BP to 1870 AD (Little Ice Age). The pattern of climate
variability in Cascade Lake generally agrees with widely recognized Holocene climate
fluctuations, though the ages differ. This may be due to a time-transgressive nature of
these climate events or, more likely, a poorly constrained age model.
Based on the interpretation of several periods of climate variability in the Ahklun
Mountains during the late Holocene, the climate of the 20th century can be placed in
context. Within the Cascade Lake record, no other period has comparable biogenic
silica) concentrations to the 20th century. This suggests an unprecedented strengthening
of the Aleutian low during this most recent period.
iv
ACKNOWLEDGMENTS
It was through the hard-work and help of many people that this project was
completed. My sincere gratitude first goes to Darrell Kaufman, as he has been a
remarkable advisor and role-model throughout every step of this project. Without his
energy during field-work, excitement over new data and encouragement during the
writing process, this thesis would lack the integrity and detail that it now has. I thank
him for always having an open office door, even when his schedule was busy, and for
asking the questions that kept me motivated.
My thanks also extend to my committee members, Jim Sample, Scott Anderson,
and Al Werner. Their comments and observations helped to tie my loose ends together.
I’ve always enjoyed our conversations both within the formal committee meeting setting
and informally in the hallway. I appreciate the extensive time and brainpower that you
poured into my thesis. I especially appreciate the help of Al Werner, because it was
through his encouragement that I choose to continue my education beyond Mount
Holyoke College.
There are many additional people who have contributed to this project with their
time, and expertise. Tom Daigle, Erin Dopfel, Mike Hamilton, Heidi Roop, Apayo
Moore, Al Werner and of course Darrell Kaufman provided much needed assistance and
ideas in the field, while Togiak National Wildlife Refuge provided the logistical support.
Feng Sheng Hu of University of Illinois and Richard Ku of University of Southern
California provided the 210Pb and 137Cs analyses and John Southon of University of
California, Irvine for the radiocarbon analysis. Jim Wittke spent a significant amount of
v
time helping analyze tephra on the electron microprobe and scanning electron
microscope. Tom Daigle, Jordan Bright, Caleb Schiff, Nick McKay and Chris Kassel
each get thanked for spending time talking over ideas with me and asking the hard
questions.
Special thanks go to my family and friends for putting up with me. Sheena
Styger, Chloe Bonamici, Christyanne Melendez, and Monica Hansen provided necessary
distractions and coffee breaks. My parents have always supported their wandering child
and continue to be the most significant role models in my life. Skip Martin has provided
me with retreats to cabins in the woods and many hugs from 3000 miles away.
This project was funded by National Science Foundation grants ARC-0455043
and ATM-0318341 awarded to Dr. Darrell Kaufman, and a Geological Society of
America graduate student research grant.
vi
TABLE OF CONTENTS
Page ABSTRACT………………………………………………………………………………ii ACKNOWLEDGMENTS………………………………………………………………..iv LIST OF TABLES………………………………………………………………...……..vii LIST OF FIGURES……………………………………………..………………………viii CHAPTERS
1. INTRODUCTION………………………………………………………………1 a. Project Objective 1 b. Background 2 c. Study Area 8
2. MODERN CLIMATE AND SEDIMENTATION AT CASCADE LAKE………9
a. Methods 9 b. Results 11 c. Discussion 13 d. Summary 15
3. GLACIAL GEOLOGY…………………………………………………………16 a. Background 16 b. Methods 18 c. Results 20 d. Discussion 23 e. Summary 26
4. CASCADE LAKE SEDIMENTS………………………………………………..26 a. Background 26 b. Methods 27 c. Results and Discussion 32
5. BIOGENIC SILICA AND CLIMATE INTERPRETATIONS…………...……..45 a. Methods 47 b. Results 48 c. Discussion 49
6. CONCLUSIONS………………………………………………………………...55
vii
LIST OF TABLES
Table Page
1. Instruments deployed at Cascade Lake………………………………….68
2. Data from sediment traps………………………………………………..69
3. Climate comparison with lichen growth-curve sites…………………….70
4. Lichen sizes and ages……………………………………………………71
5. Equilibrium line altitudes for glaciers…………………………………...72
6. Lake cores recovered for this study……………………………………...73
7. Tephra descriptions from Core 02……………………………………….74
8. BSi and meteorological correlations……………………………………..75
viii
LIST OF FIGURES
Page
1. Summary of Holocene climate fluctuations………………………………... 76
2. Map of Alaska and the Ahklun Mountains………………………………… 77
3. Color-infrared aerial photograph of the study area……………………….... 78
4. Types of instruments deployed for this study……………………………… 79
5. Data recovered from 2004-2005, 2005-2006 sediment traps………………. 80
6. Air and water temperature data from the study area……………………….. 81
7. Relationship between Dillingham, Bethel and Cascade temperatures…….. 82
8. Cascade glaciers…………………………………………………………… 83
9. Equilibrium-line-altitude curves…………………………………………… 84
10. Moraines of Cascade Glacier and the East Inflow cirque………………….. 85
11. Lichen growth curves………………………………………………………. 86
12. Retreat rates of Cascade Glacier……………………………………………. 87
13. Bathymetric map…………………………………………………………… 88
14. Methods used in lamination counting……………………………………… 89
15. Acoustic profile from Core 02 vicinity…………………………………….. 90
16. Lithostratigraphic logs…………………………………………………….. 91
17. Correlations with long and surface………………………………………… 92
18. Test pit wall………………………………………………………………… 93
19. 210Pb and 137Cs profiles…………………………………………………….. 94
20. Geochemical data from the 2.85 m tephra and Aniakchak reference samples. 95
21. Scanning Electron Microscope images of tephra shards………………….. 96
22. Age model…………………………………………………………………. 97
23. Magnetic susceptibility…………………………………………………….. 98
24. Organic matter content and bulk density…………………………………... 99
25. Lamination thickness……………………………………………………… 100
26. Biogenic silica…………………………………………………………….. 101
27. Biogenic silica of the last 2200 years……………………………………… 102
28. Core 02 proxy data by age…………………………………………………. 103
1
CHAPTER 1:
INTRODUCTION
The high latitudes have been shown to be particularly sensitive to global climate
change, with dramatic environmental changes induced by subtle climate variations
(Overpeck et al., 1997). Throughout the Quaternary, inter-annual and century-scale
climate variation has been the norm; however, the timing, synchronicity and forcing
mechanisms of these fluctuations are still poorly understood and necessitate further work.
The investigation and interpretation of these past environmental changes provide a
context within which to place our understanding of present and future climate
fluctuations.
Lake sediment cores have proven particularly successful in providing valuable
interpretations of Holocene climate fluctuations (Pienitz et al., 2004). Lake sediments
provide continuous records of glacial fluctuations, lake productivity, sediment fluxes and
a wide variety of other environmental changes induced by climate variability. Using
proxy indicators available in lacustrine sediments, it is possible to determine the range of
non-anthropogenic Holocene climate variability. The naturally induced events of the
Holocene (e.g., Neoglaciation, Medieval Warm Period, and Little Ice Age) can then be
compared with the twentieth-century climate changes that are related to anthropogenic
forcing.
Project Objective
The overarching goal of this study is to determine the range of climate variability
preserved in the sediments of Cascade Lake, Alaska. This is accomplished by: (1)
2
establishing an understanding of the modern and historic climate and sedimentation
patterns at Cascade Lake, (2) determining the recent glacial fluctuations within the
drainage basin of Cascade Lake, (3) identifying the major sedimentation patterns and
fluctuations in Cascade Lake sediment cores, and (4) using proxy indicators to interpret
Holocene climate variations.
Background
Holocene Climate Variability. Initial paleoclimate reconstructions from the marine and
ice cores suggested that the Holocene was a period of relative stability compared to the
large-scale climate fluctuations of the Last Glacial Maximum. Modern alpine glaciers
were originally assumed to be merely the remnants of maximum Pleistocene ice extent.
However, it is now recognized that this ‘quiescent’ period includes a series of
pronounced climate fluctuations (e.g., Mayewski et al., 2004). Multiple cooling events
have been identified that punctuate the current interglacial period; perhaps the most well
known is the 400-year-long Little Ice Age (LIA). This period of multiple “glacial
expansion(s) subsequent to the maximum Hypsithermal shrinkage” has been termed
Neoglaciation (Porter and Denton, 1967).
One of the first studies to systematically approach North American climate
fluctuations of the late Holocene was by Denton and Karlén (1973). Their study of the
expansion and contraction of 51 North American and European alpine glaciers
recognized three prominent intervals of glacial expansion: 3300-2400, 1250-1050 cal yr
BP (where “BP” = before 1950 AD), and the Little Ice Age (1500-1900 AD). This
millennial-scale pattern (~1500 and 2000-2800 year climate periodicity) of widespread
3
late Holocene cooling events is gaining support as evidence of similar climate variability
emerges from other high-resolution records (Bond and Showers, 1997; Campbell et al.,
1998; Bianchi and McCave, 1999; Leonard and Reasoner, 1999; Noren et al., 2002;
Mayewski et al., 2004). Mayewski et al. (2004) summarized this supporting evidence in
a study of ~50 spatially distributed paleoclimate records. They determined that a
minimum of six global Holocene cooling events are associated with this pattern occurring
at 9000-8000, 6000-5000, 4200-3800, 3500-2500, 1200-1000 and 600-150 cal yr BP
(Figure 1).
The millennial-scale climate variability of the Holocene is not only marked by
cool periods. Mild, warm and dry periods have also been recognized at 10600-9300,
7900-6300, 2700-1500, 960-610 cal yr BP (Figure 1, O’Brien et al., 1995). The warm
period of approximately 1000 cal yr BP (1000-1300 AD) has been termed the Medieval
Warm Period (MWP). This period is often cited as being similar to the modern 20th
century warming; however, a compilation by Crowley and Lowrey (2000) determined
that the MWP was only ~0.20°C warmer than the LIA. The spatial and temporal patterns
and the amplitude of this warm period must be resolved before it can be used as a model
for 20th century warming.
A forcing mechanism for this millennial-scale climate periodicity has yet to be
determined. Some evidence suggests that these cycles are related to Dansgaard-Oeschger
like ice-rafting events, and therefore linked to North Atlantic thermohaline circulation
and a strongly coupled ocean-atmosphere system (Alley et al., 1997; Bond and Showers,
1997). Other evidence suggests that external factors, such as persistent solar variability,
4
explain the widespread distribution of these climate events (Denton and Karlén, 1973;
Mayewski et al., 1997; Bond et al., 2001).
Neoglaciation in Alaska. The glacial records of past climate variability in Alaska
generally concur with the more widely recognized events of the late Holocene. Calkin’s
(1988) summary of Holocene glaciation in Alaska focused on glacial moraines that
indicate four major periods of glacial advance including: an initial advance between
6000-5000 cal yr BP, two distinct advances at 3600-2000, and 1700-1100 cal yr BP and
the most recent advance of the Little Ice Age at 750-400 cal yr BP. While these periods
of glacial advance have been recognized throughout Alaska, their timing, duration and
magnitude vary considerably. For instance, the maximum Holocene ice extent in regions
of the Kenai Peninsula was achieved during the 3600-2000 cal yr BP advance whereas
many other local areas experienced maximum ice extent during the LIA (Calkin, 1988).
Some of the variation in the timing of Neoglacial advances could be attributed to
the difficulty in identifying geomorphic evidence for the extent of early Neoglacial
advances, which might have been obliterated by more extensive LIA advances.
However, supporting evidence has been derived from proxy records such as pollen, lake-
level and tree-ring studies. In southern Alaska these supporting studies indicate that
periods of glacial advance identified in glacial geomorphic studies correspond to periods
of increased wetness (e.g., Hansen and Engstrom, 1996; Abbott et al., 2000; Calkin et al.,
2001).
Pacific Ocean and Alaska Climate. Interdecadal climate variability in the North Pacific
region has been well documented throughout the last century and likely influences the
5
millennial-scale climate fluctuations of the Holocene ( Minobe and Mantua, 1999;
Schneider and Cornuelle, 2006). Although the mechanism for this interdecadal climate
variability is unknown, a basin-wide phenomenon that originates in the tropics and is
closely related to the shorter-term oscillations of sea-surface temperature and sea-level
pressure has been suggested (D’Arrigo et al., 2005). The primary two short-term
oscillations are the El Niño-Southern Oscillation (ENSO) and the Aleutian low-pressure
cell (AL). ENSO regimes typically persist for 6-18 months and are responsible for shifts
in the locations of warm tropical waters. These shifts in sea-surface temperatures (SST)
in turn affect the regional precipitation and mean annual temperatures (Papineau, 2001;
Neal et al., 2002). The dominant mode of North Pacific wintertime atmospheric
circulation is the Aleutian low. A deepened AL results in a strengthened southerly flow
over the eastern Pacific, and warmer temperatures along the western coast of North
America (Rodionov et al., 2005). Each of these components contributes to the Pacific
Decadal Oscillation (PDO; measured by SST), which is one of the best summaries of
North Pacific interdecadal climate variability.
The PDO has its primary signature in the North Pacific Ocean with a secondary
imprint in the tropics (where the ENSO fingerprint is primary) and a periodicity of
approximately every 20-30 yr. Over the historic period the PDO has had several
documented shifts in its primary modes (Kayano et al., 2005). ‘Cool’ phases (cold waters
in the Gulf of Alaska) persisted from 1900-1924 and 1947-1976. Subsequent ‘warm’
phases occurred from 1925-1946 and 1977 to the mid-1990s. These interdecadal shifts in
the PDO have a strong link to interannual variation of the AL. Warm modes of the PDO
are associated with a general strengthening of the AL; however, the interannual variance
6
of the AL during the warm PDO phases is two to three times greater than other periods
(Minobe and Mantua, 1999). Though these patterns of variability in the closely coupled
ocean-atmosphere system of the North Pacific have only been documented over the last
100 yr, paleoclimate studies suggest that these patterns have persistently influenced
southern Alaska climate throughout the Holocene. Generally, the association between
increased wetness and periods of glacial advances in northwestern North America implies
a strong influence of the AL on Holocene glaciation (Spooner et al., 2003; Anderson et
al., 2005).
Previous Work in Ahklun Mountains, Alaska. Although significant paleoclimate research
has been completed in Alaska, relatively little attention has focused on the late Holocene
in the southwestern portion of the state. Previous work in southwestern Alaska has
largely focused on the Ahklun Mountains. The Ahklun Mountains are a 250-km-long
mountain belt trending southwest to northeast between Bristol Bay and Kuskokwim Bay
lowlands (Figure 2). They form the largest mountain range west of the Alaska Range and
harbored a large ice cap during the late Wisconsin that was separate from the Cordilleran
Ice Sheet. Most of the paleoclimate research completed in the region has focused on
Pleistocene and Younger Dryas glacial advances (Briner and Kaufman, 2000; Kaufman et
al., 2001; Manley et al., 2001; Briner et al., 2002; Hu and Shemesh, 2003; Hu et al.,
2006).
Studies of the Holocene climate in the Ahklun Mountains remain rather limited.
A lacustrine sediment sequence recovered from Arolik Lake in the southwestern Ahklun
Mountains produced a 33,000 yr record of climate variability (Kaufman et al., 2003).
While high climate variability was observed in the Arolik Lake record, sampling
7
resolution was too coarse to resolve Holocene fluctuations. A more detailed study (Hu et
al., 2003) observed regular oscillations in lake productivity coinciding with solar
variability, but this study focused on a portion of the core ranging in age from 12,000 to
2300 cal yr BP and did not address Neoglacial cooling events.
Studies that directly address late Holocene paleoclimate in the Ahklun Mountains
include a pollen study by Axford and Kaufman (2004) and a combined lacustrine/glacial-
moraine study by Levy et al. (2004). While Little Swift Lake, in the northwestern
Ahklun Mountains, does not contain clear evidence of specific Neoglacial cooling events,
the pollen record does indicate progressive climatic deterioration (cooling) throughout the
late Holocene (Axford and Kaufman, 2004). Levy et al.’s (2004) work on Waskey Lake
is the most complete work to date on Ahklun Mountain Neoglaciation. They identified
multiple late Holocene glacial advances by the presence of several discrete end moraines
in the forefields of the modern glaciers. Both the geomorphic evidence and the lake-
sediment record indicate that the peak Neoglacial activity occurred at Waskey Lake
between 700 and 200 cal yr BP. The lake sediments indicate an onset of Neoglaciation at
3100 cal yr BP. This age is significantly younger than the ages suggested for major
Neoglacial advances across the rest of Alaska (Brooks Range, Kenai Mountains, and
Seward Peninsula). This suggests that the Ahklun Mountains may be slightly out of
phase with widespread climate events of the late Holocene. Without additional records of
Neoglaciation in the Ahklun Mountains, however, it is difficult to determine the full
significance of the Waskey Lake record. If the timing of glacial activity at Waskey Lake
matches other places across the Ahklun Mountains this would imply that an unknown
regional-scale climate forcing exerts a stronger influence than the large-widespread
8
climate forcing affecting other areas of Alaska and potentially the globe. Cascade Lake is
located only 38 km from Waskey Lake providing an opportunity to test the
reproducibility of the Waskey Lake Holocene climate record reported by Levy et al.
(2004).
Study Area
Presently, over 100 active glaciers are located in the northern portions of the
Ahklun Mountains where numerous peaks are over 1500 m a.s.l. The largest of these
northern glacial complexes are Chikuminuk Glacier (~5.6 km2) and an unnamed glacier
(hereafter, “Cascade Glacier”, ~3.4 km2) to the west and south respectively (Figure 3).
Cascade Lake (60° 8’ 11.9’’N; 159° 23’ 57.3’’W) lies 5.4 km downstream from the
primary outflow of Cascade Glacier at 335 m elevation.
The drainage basin area of Cascade Lake is 43.7 km2, 12% of which is occupied
by active glaciers. The area is vegetated with stands of alder and willow shrub and upland
tundra. Cascade Lake (3.1 km2) has multiple inflows, both melt-water streams and non-
glacial channels, though the primary sediment contributor is Cascade Glacier (Figure 3).
Cascade Lake is formed in late Wisconsin drift underlain by Jurassic volcaniclastic
sandstones, conglomerates and argillites that are members of the Gemuk group (Box et
al., 1993). Two interpreted faults bisect the field area, one crossing the northern portion
of the lake and the other crossing mid-way down the melt-water channel of Cascade
Glacier (Hoare and Coonrad, 1959).
9
CHAPTER 2: MODERN CLIMATE AND SEDIMENTATION AT CASCADE LAKE
In order to make quantitative interpretations of Holocene climate variations
preserved in lake sediments, it is necessary to understand the climate factors affecting
sediment transfer and deposition. Factors such as temperature, rainfall, and snowmelt
intensity, have a direct impact on the sediment characteristics preserved in lakes (Leeman
and Niessen, 1994; Hardy et al., 1996; Noon et al., 2001; Francus et al., 2002). A small-
scale study of the modern sedimentation and climate at Cascade Lake was undertaken to
enable a better understanding of the influence of climate on its sediment record and to
evaluate the potential for annual sedimentation patterns.
Methods
Moorings. Instrument moorings were deployed at water depths of 24.2 m (mooring 1)
and 52.4 m (mooring 2) at glacial distal and proximal sites, respectively, within Cascade
Lake (Figure 3, Table 1). Each mooring consisted of a rock anchor, two sediment traps,
two temperature loggers and a buoy. The sediment traps were made from 12.5-cm-
diameter funnels that fed rigid lexan tubes with an inner diameter of 1.6 cm (focusing
factor = 62 times) (Figure 4). Traps were placed at ~1 m above the lake bottom and ~3 m
below the lake surface. Buoys were set at ~2 m below the lake surface to avoid exposure
during lower winter lake levels or freezing in winter ice. Each sediment trap was paired
with an Onset Water Temp Pro temperature logger that recorded temperature every hour
beginning June 28, 2004.
10
Trapped Sediments. The sediment traps were recovered on August 10 and 11, 2005 and
on July 19th, 2006. The sediment was described and sub-sampled into stratigraphically
distinct units visible in the receiving tube. The sediment from the 2005 traps at mooring
2 was analyzed for bulk-sediment density, and 2005 sediment from mooring 1 was
processed for organic-matter content (OM), grain size, and biogenic silica content (BSi).
The 2006 sediment from mooring 2 was analyzed for bulk-sediment density and BSi
content.
Sediment grain-size distribution was determined using 0.25 cm3 samples analyzed
by a Coulter LS230 laser diffraction analyzer. Samples were pre-treated with 30% H2O2
to remove organics and 1M NaOH to dissolve BSi, and stored in sodium
hexametaphosphate (a dispersing agent) to prevent flocculation prior to analysis. Each
sample was analyzed four times for 80 seconds and all runs were averaged for the
samples final grain-size estimate. OM content was determined by percent loss-on-
ignition (LOI) and BSi analysis was completed using the methods of Mortlock and
Froelich (1989, see Lake Sediments Methods).
Air Temperature. Onset Hobo temperature loggers were installed on the western shore of
Cascade Lake and ~0.6 km from the modern Cascade Glacier front (Figure 3). The
lakeshore logger was placed inside a solar radiation shield and mounted at a height of ~3
m on a large willow tree in a stand of alder (Figure 4). The logger recorded air
temperature every hour for a period of June 28, 2004 to August 9, 2005, and from August
12, 2005 to July 20, 2006. The glacier front air temperature logger was mounted on
stabilized copper poles ~2 m above the ground and recorded temperature every hour from
August 13, 2005 to July 21, 2006.
11
Results and Discussion
Sediment Traps. Mooring 1 was recovered from its deployed location in both 2005 and
2006; mooring 2 was recovered from deeper water than which it was originally deployed
in 2005. The mooring likely moved (drifted or was carried by lake ice) during the
collection year, but the sediment traps and temperature loggers were all intact and
seemingly undisturbed. Upon recovery the top traps were covered in a thick algal mat
and had collected limited porous sediments with live organisms. The bottom traps
contained significantly more sediment than the top traps (2005 mooring 2 had slightly
overflowed into the collection tube) and were sub-sampled into an average of five
stratigraphic units (Figure 5). Sediment in the bottom traps was typically gray to gray-
green with organic material overlaying dense mud with gradational contacts between
units.
Significantly less sediment accumulated during 2005-2006 than in 2004-2005.
Unfortunately, a fingerling fish was caught in the receiving tube of 2006 mooring 1 and
made field collection of the sediment impossible. However, mooring 2 collected 8.5 cm
of sediment in 2006 compared to the 18.1 cm in 2005. The average sediment
accumulation in the bottom trap of mooring 2 was 0.10 g/cm2/yr in 2004-2005, and 0.03
g/cm2/yr in 2005-2006, a 70% reduction (Tables 1 and 2). Because of possible
complications caused by the movement of mooring 2 during the year and the overflowed
collection tube, I focused on 2004-2005 mooring 1 for detailed analysis. Each of the
stratigraphic units was analyzed for OM, grain size, and BSi content. The 2006 mooring
2 sediments were also analyzed for bulk density and BSi content for comparison (Figure
12
5). OM, grain size, and BSi content all covaried in the sediment from mooring 1 in 2005,
with their lowest values in the middle units (Figure 5). The sediment from mooring 2 in
2006 contained less BSi overall (16% reduction in BSi percentage) and did not show
decreased BSi content in the center units. Adjusting BSi values for the dilution caused by
the increased sedimentation in the 2004-2005 collection year, results in a 77% reduction
in BSi flux during 2005-2006 (Table 2).
The results from the bottom trap of 2004 mooring 1 exhibit a seasonal pattern
(Figure 5). Because the traps were not equipped to record the timing of sediment
deposition within the collection tubes, the seasonality of each lithologic unit is not
known. Nonetheless, the decrease in OM, BSi, and grain size in the central part of the
receiving tube of mooring 1 is indicative of winter sediment deposition in a glacial-fed
lake. The increase in BSi in the uppermost sediments (units 1 and 2) and the coarse
sediments (coarse silt) in unit 1 might indicate higher stream discharge associated with
spring melt-out. Mooring 1 was located farther from the glacial melt-water inflows, and
therefore was less sensitive to the glacial melt-water influences and may not have
captured the seasonal signal as well as the more proximal site might have.
The 70% decrease in sediment flux and the 77% reduction in BSi deposition
between the two collection years are indicative of a strong influence of interannual
weather differences on sediment transport and accumulation. Spring and summer (May
to September) surface-water temperatures (air temperature not available for 2005) were
~1.5 °C colder in 2006 than in 2005, and in July of 2006, significantly more snow
remained on the landscape. Additionally, less (no quantitative measurements exist) melt-
water occupied the channels of the Cascade Glacier outwash plain than in the previous
13
year. There is no precipitation data for Dillingham post-1998; however, precipitation
data from Bethel over the collection years shows a 7.6 cm decrease in annual
precipitation from 2005 to 2006 (62.5 vs. 54.9 cm; WRCC, 2006). Possibly the main
pulse of sediment associated with peak melt-water discharge had not occurred by the time
the traps were recovered in July 2006. More likely, the cooler spring/early summer
temperatures and decreased precipitation retarded melt-water production (i.e. peak
discharge) and decreased the sediment flux and associated nutrients for algal growth.
Water Temperature. The four temperature loggers were recovered in August 2005 and
July 2006 (Figure 6). The water temperature data indicate that Cascade Lake is dimictic
with fall mixing occurring in late October and spring melt-out and mixing occurring in
late May to early June. Thermal mixing during the fall lasted 21 days in 2004 and 13 days
in 2005 whereas spring mixing took place over 8 days in 2005 and 4 days in 2006. Fall
cooling and spring warming leading to or following overturning were gradual at a rate of
about 1 °C per 9 days. The lake-surface temperatures are similar at both mooring sites
(0.1 °C warmer at the mooring 1 site). During the ice-free months, the lake-surface
temperatures track air temperature with a minimal lag time (hours) and average 5 ± 9 °C
cooler (based on hourly averages).
Air Temperature. In 2005, the lakeshore temperature logger malfunctioned and stopped
recording on December 22, 2004 (after 176 days, Figure 6). A new lakeshore logger and
the glacier terminus logger were deployed in August 2005 and downloaded in July 2006.
The radiation shield at the glacier terminus was knocked off its mount and the logger
14
recorded a stable temperature of ~-7 °C from the end of January to the beginning of June
2006, probably because the fallen logger was buried by snow during this interval.
Freezing air temperatures at the lake were first reached in September and
generally remained below freezing by November in both 2004-2005 and 2005-2006.
Temperatures averaged above freezing starting in early May. For the period successfully
recorded by both lakeshore air temperature loggers (August 12-January 22), the average
2004 temperature was 1.2 °C and the average 2005 temperature was –1.4 °C. The coldest
recorded temperature in 2004 was –28 °C in early December, and in 2006 it was –40 °C
in late January. The onset of freezing air temperatures and the timing of above-freezing
temperatures in the spring correspond well to the build-up and melt-out of lake ice, just
prior to and after lake overturning. Air temperatures lead lake ice-up and melt-out by
about a month.
The glacier logger first recorded freezing temperatures two months earlier than
the lake, and minimum temperatures occurred just before it went isothermal (-35°C,
Figure 6). During the summer (JJA) the lakeshore air temperature averaged 2 ± 1 °C
warmer than the glacier, whereas the lakeshore was 2 ± 4 °C cooler during the winter
(DJF). With an elevation difference of 340 m these two temperature stations record an
environmental lapse rate of ~6 °C/km during the summer and an inverted profile in the
winter.
The air temperature logged at Cascade Lake can be compared with other climate
stations in the region. The goal of this comparison was to determine whether the longer-
term temperature records of the climate stations could confidently be extrapolated to
Cascade Lake. The two closest climate stations are Bethel (~148 km to the northwest)
15
and Dillingham (~140 km to the southeast), where daily maximum and minimum
temperature data are available (PNWPEST, 2006). I compared the daily maximum,
minimum, and average daily temperatures from the climate stations (mean of two values)
and from the Cascade Lake hourly air-temperature logger (mean of 24 values) and found
the strongest correlation between the average daily temperature (r = 0.97 and 0.96 for
Dillingham and Bethel, respectively). Daily temperature at Cascade Lake averaged 3.5 ±
3.1 °C (n = 498) cooler than Dillingham and 1.1 ± 3.4 °C cooler than Bethel (Figure 7).
The temperature differences between Cascade Lake shoreline and Dillingham vary by
season. During the summer, Cascade Lake was 2 °C cooler than Dillingham and 5 °C
colder in the winter (Figure 7). The temperature differences between Dillingham and
Cascade Lake indicate a summer environmental lapse rate of 7 °C/km. This is higher
than the typical lapse rates of 6 °C/km (Porter, 2001), but similar to that of the lake and
glacier temperature loggers. The somewhat higher lapse rate may be due to the more
southerly and coastal location of Dillingham. Although temperature at both Dillingham
and Bethel are strongly correlated with Cascade Lake, I chose to focus on the climate
record at Dillingham because the temperature record there extends back to 1919,
(compared with 1949 at Bethel).
16
CHAPTER 3: GLACIAL GEOLOGY
Glaciers are sensitive to climate variations and are known to have fluctuated
during the late Holocene in mountains across Alaska (Calkin, 1988; Calkin et al., 2001).
The study of glacial deposits can provide information on the timing and magnitude of
climate change that caused their deposition. By combining a glacial-geologic and a lake-
sediment study from the same drainage, the two complimentary records of climate
variability can be compared.
Background
Lichenometry. Lichenometry is a useful dating method applied to late Holocene
landforms (Locke et al., 1979). In many cases, the historic record is too fragmented,
carbonaceous material is absent or the landforms are too young (<500 yr) for reliable
radiocarbon dating, or the tree-ring record is too short. Lichenometry is based on the
premise that lichens grow at a predictable rate once they colonize a surface. Rhizocarpon
geographicum and R. alpicola are the two most commonly used species because they
have slow growth rates (0.02-2.0 mm/yr; Innes, 1985). These two species have slightly
different growth rates and they are impossible to distinguish in the field (Werner, 1990).
Most authors report data as Rhizocarpon geographicum senso lato (s.l.) and assume that
this includes measurements of both R. geographicum and R. alpicola thalli.
The growth rate of Rhizocarpon is dependent on a variety of factors including, but
not limited to: substrate lithology, moisture availability, temperature, altitude, and
17
snowcover (Innes, 1985). Each of these factors can vary significantly from one location
to another, therefore, it is important to use a growth curve developed either from the
region of interest, or from a similar environment. Lichen ages reflect the timing of
landform stabilization and because moraines take some time to stabilize following the
retreat of abutting glacial ice, they provide minimum ages for moraine deposition.
Additionally, actively eroding landforms can not produce reliable lichenometric ages and
should not be used (O’Neal, 2006).
Equilibrium-Line Altitude. The equilibrium-line altitude (ELA) serves as one of the most
useful parameters to quantify the effect of climate on a glacier. The ELA separates the
areas of accumulation and ablation and is controlled by a variety of factors including
glacier hypsometry, glacier aspect, and other local climate conditions (Torsnes et al.,
1993). Changes in ELA over time are generally most sensitive to winter precipitation
and summer temperature, thereby enabling effective paleoclimate reconstructions (Dahl
and Nesje, 1992).
A variety of methods have been developed to reconstruct former ELAs. Torsnes
et al. (1993) compared various methodologies and determined that the accumulation area
ratio (AAR) method resulted in the least error when calculating the ELA of modern
glaciers with ELAs known through mass-balance measurements. The accumulation area
ratio is defined as the ratio of the accumulation area to the total glacier area and is used to
determine steady-state ELA. Meier and Post (1962) demonstrated that the AAR for
steady-state glaciers in northwestern North America ranged from 0.5 to 0.8. The average
of these values (0.65) is often used in ELA reconstructions (Porter, 2001).
18
Methods
Lichenometry. Glacial and fluvial deposits believed to be Holocene in age were identified
on aerial photos and then studied in the field, or in some cases deposits were first
recognized in the field. These deposits included moraines of the most accessible glaciers
in the study area, fluvial terraces in the outwash plains of these glaciers, and stable
surfaces that were formed during glacier retreat. The long and short axes of large
Rhizocarpon geographicum s.l. thalli were measured using digital calipers. Landform
surfaces were searched until no larger thalli could be found. This involved searching the
entire landform surface and measuring approximately 15 lichen thalli. Generally, the
largest lichens were found on stable slopes (crests of moraines) and in protected
depressions on landform surfaces. Lichens with a long axis that measured over 20%
larger than the short axis were rejected because of this non-regular growth pattern (Calkin
and Ellis, 1980).
Because a lichen growth curve has not been developed for the Ahklun Mountains,
the lichen sizes were compared to the growth rates at five other Alaska locales including:
the Brooks Range, Seward Peninsula, central Alaska Range, Wrangell-St. Elias
Mountains, and southern Kenai Mountains (Figure 2). Solomina and Calkin (2003)
recently revised the growth curves for each of these areas based on additional control
points, newly calibrated 14C ages, and standardized fitting routines. These growth curves
either used the single largest thallus diameter or the averages of the five largest
diameters. Both of these values were calculated for the Cascade River valley data and
compared with the corresponding growth curves. Uncertainty in the derived ages results
19
from the variety of climates represented by the five growth rates and their dissimilarities
from the Cascade River valley and measurement errors. Of the five areas for which
Solomina and Calkin (2003) developed growth curves, the mean annual and summer
temperature (and likely the precipitation) of the Cascade Lake area are most similar to
Seward Peninsula and therefore was used for age assignments (Table 3).
Equilibrium-Line Altitudes. Eight glaciers were chosen for ELA analysis, including those
visited during the 2005 field season, and those associated with laterally extensive and
prominent moraines on the 1957 aerial photographs. These eight glaciers include the
Cascade and Chikuminuk valley glaciers and six cirque glaciers (Hungry Joe Cirque,
Divided Cirque, East Inflow Cirque, Cirque 1, Cirque 2, Cirque 3, Figure 8). Four of the
glaciers are within the Cascade Lake drainage basin and all are located within 9 km of the
lake.
I used the AAR method for ELA reconstructions with an AAR value of 0.6 to
facilitate direct comparison with the Waskey Lake area (Levy et al., 2004). The extent of
ice in the 1957 aerial photographs and the reconstructed LIA ice limits (based on field
glacial-geomorphic evidence and aerial photograph interpretation) were plotted on a
1:63,360-scale topographic map (1973) in ArcGIS v. 9 (i.e., ‘modern’ = 1957). This
assumes that the 1973 contours were a reasonably accurate representation of the 1957
glacier surface. The area of each glacier was measured at ~100-ft-altitude intervals for
both the modern and reconstructed LIA ice limits. A multi-order polynomial was fit to
the data and used to determine the altitude corresponding to an AAR of 0.6 (40% of the
total area lies within the zone of ablation; Figure 9).
20
Results and Discussion
Lichenometry. A total of 160 lichens with circular thalli were measured on 13 deposits in
the forefield of five extant glaciers (Figure 8, Table 4). The outer moraines represent the
largest Neoglacial advance while other deposits further up-valley represent smaller
subsequent readvances, delays in retreat, or are deposits of a non-glacial origin (debris-
flow deposits, river terraces, etc.). Of the five forefields, only two contained stable late
Holocene moraines with measurable lichens, the other deposits were either unstable
(slumped) over this period, or could not be clearly associated with the maximum ice
advance . Lichens were measured on two distinct moraine deposits on both the left and
right lateral sides of Cascade River valley and the terminal moraine of the East Inflow
cirque (Figures 8, 10). The age estimates for these five moraines derived from the range
of growth rates span from 860 cal yr BP based on the slow growth curve (Brooks Range)
to 20 cal yr BP using the fast growth curve (Kenai Peninsula, Figure 11). The moderate
growth curve from the Seward Peninsula (with a ~30-yr colonization time, see below)
yields an average age of 90 cal yr BP (~1860 AD) for these LIA maximum moraines.
Although the control points from the Seward Peninsula growth curve are not well-
constrained around this time period, the Alaska Range growth curve yields a similar age
for this lichen size. Because lichens date the stabilization rather than the initial
deposition of a moraine, this age is a minimum estimate of the LIA maximum in the
Cascade River valley.
Two of the lichen stations (12, 13) are up-valley of the LIA terminal moraine of
the Cascade Glacier and were used to calculate a retreat rate. Lichens were measured on
stable boulders overlying bedrock that were deposited upon ice retreat from the site.
21
Station 12 is located 990 m from the 2004 ice margin and was dated to 1945 AD using
the Seward Peninsula growth curve (with the 17-yr colonization time of Solomina and
Calkin (2003); Table 4). Station 13 is 540 m from the terminal moraine with an age of
1965 AD, but this age is too young as the site lies ~30 m beyond the 1957 ice front. This
implies that the 17-yr colonization time built into the Seward Peninsula growth curve
may be too short for the Cascade Lake area. A longer colonization time of about 30-yr is
also indicated by the lack of lichens on surfaces exposed after 1984 and by only the
earliest stages (<1 mm) of lichen colonization on the surfaces exposed between 1957 and
1984. The lichen ages reported in this thesis are therefore adjusted to be ten years
younger than the ages derived from the Seward Peninsula growth curve.
The rate of retreat at Cascade Glacier was reconstructed by constraining the ice
margin of Cascade Glacier at five points, using lichen ages and aerial photographs,
namely: the 2004 observed ice front, 1984 and 1957 aerial photographs, lichen station 12
with an age of 1935, and the maximum LIA moraine dated to ~1860 AD based on
lichenometry. These points indicate that Cascade Glacier retreated at the following
average rates: 1860-1935 = 5 m/yr; 1935-1957 = 21 m/yr; 1957-1984 = 8 m/yr; 1984-
2004 = 16 m/yr (Figure 12). The overall retreat rate for Cascade Glacier since the LIA
maximum was 11 m/yr.
The retreat rate of Cascade Glacier from its maximum LIA extent was not steady
over the last 140 yr, reflecting climatic variations and topographic constraints. The
lowest retreat rate was documented between 1860 and 1935 (5 m/yr). This time period
includes the climate recovery following the termination of the LIA. The shift to a rapid
retreat between 1935 and 1957 (21 m/yr) is more difficult to explain. The period of
22
1947-1976 coincides with a ‘cool’ PDO and weakened AL, which would have produced
cooler conditions in southwest Alaska (Papineau, 2001). These conditions should have
inhibited glacier retreat, whereas the warmer temperatures associated with the preceding
‘warm’ PDO (strengthened AL) period of 1925-1946 should have promoted glacier
retreat. Glaciers respond more rapidly to the affects of summer ablation on the glacier
toe than they do to winter accumulation during positive mass-balance years. Because the
response time of mountain glaciers averages 10-15 yr (Porter, 1986), the retreat rate of
Cascade Glacier between 1946 and 1957 may be out-of-phase with the PDO and AL and
may not have yet responded to the positive mass-balance years of 1947-1976. The 1957
and 1984 aerial photographs show the terminus of Cascade Glacier at nearly the same
location in a bedrock-constrained, narrow section of the valley, which explains the lower
retreat rate (8 m/yr) during this period. The most recent shift (post-1984) in retreat rates
to 16 m/yr is likely the glacier response to ‘warm’ PDO shift from 1977 to 1995 with
some lag in response time.
The multiple late Holocene advances recognized by Levy et al. (2004) at Waskey
Lake (~30 km south of Cascade Lake) contrast with the single LIA moraines recognized
in Cascade Lake study area. Both this study and Levy et al.’s (2004) used the same
lichenometric methods, yet Levy et al. (2004) identified three periods of ice advance at
900-1800, 350-1050, and 200-650 cal yr BP, none of which are represented in this study
area. Conversely, the most extensive advance in my study area took place at 90 cal yr
BP, but is not represented in Waskey Lake valley. This discrepancy in Neoglacial ice
advance is difficult to understand. The glaciers of both study areas are similar in size and
occupy mountains with similar physiography and climate regime. None of the glaciers
23
overlie unconsolidated sediment that may lead to variations in sliding velocities; rather,
they directly overlie bedrock. The Waskey Glacier is underlain by granodiorite, whereas
the Cascade River valley glaciers overlie meta-sedimentary rocks with significant
regional faulting. It is a possibility that differences in sub-glacial hydrology (bedrock
permeability) may cause some of the variations in ice extent between the Cascade River
valley glaciers and Waskey Glacier.
The late 19th century age for the LIA moraines in Cascade River valley
corresponds to LIA moraine ages elsewhere in Alaska. In the northeastern Brooks
Range, for example, LIA retreat occurred around 1890 AD (Evinson et al., 1996). Late
19th century moraine stabilization has also been recognized in the southeastern coastal
glaciers, and the Kenai Mountains (Wiles and Calkin, 1994; Calkin et al., 2001). Taken
together, these ages suggest that the second half of the 19th century was marked by
widespread moraine formation across Alaska.
Equilibrium-Line Altitude. The modern glaciers (1957) in the area of Cascade Lake have
an average size of 1.6 km2 (n = 8) and an average ELA of 980 m ± 90 m (Table 5). The
glaciers in the study area extended 0.2 to 1.7 km farther down-valley during the LIA, an
average of 0.6 km (65%) larger than their 1957 counterparts. Chikuminuk Glacier
experienced the largest down-valley expansion of 1.7 km. Cascade Glacier’s ELA of 980
m coincides with the mean of the average ELA and its 1.4 km down-valley expansion
during the LIA is the second largest.
The average ∆ELA (lowering of ELA from LIA maximum to 1957) for the
Cascade Lake area was 23 ± 13 m (Table 5). ∆ELA ranges from 45 m (Cirque 2) to 5 m
24
(Cirque 1). The difference in ∆ELA between these two cirque glaciers is unexpected
because they have similar southerly aspects, are within 1 km of one another, and have
similar modern ELAs (1010 and 1050 m respectively), but Cirque 1 contains the most
exposed glacier with an extensive, low-relief forefield, which may explain its limited
expansion during the LIA. ∆ELA does not appear to covary with glacier area, head
altitude, or aspect (Table 5).
The modern ELA of 980 ± 90 m determined in this study is similar to a GIS
assessment of 106 cirque glaciers in the Ahklun Mountains completed by Manley (1999).
Manley estimated a regional ELA of 929 ± 127 m, with 73% of the variance in ELA
accounted for by microclimate variables (winter precipitation, summer temperature,
aspect, etc.). Manley (1999) also demonstrated that the Ahklun Mountains cirque
glaciers are sensitive to climate changes that affect accumulation and ablation. For
example, a 50 m rise in the ELA would result in a 40% decrease in the total glacier area.
The only other ELA analysis from southwestern Alaska is that of Levy et al. (2004) for
Waskey Lake. Their ∆ELA estimate for the LIA (35 ± 22 m) overlaps with the Cascade
Lake study area. Additionally, Skiorski (2004) and Daigle and Kaufman (2006)
determined ∆ELA’s of 25 ± 17 m in the Brooks Range and 50 ± 20 m on the Kenai
Peninsula, respectively. In contrast, ∆ELA values reported for the LIA in Alaska by
Calkin et al. (1985) range from 100-200 m. Their calculation of ‘modern’ ELA was
based on mass-balance data from 1976-1982, whereas Levy et al. (2004) Skiorski (2004),
Daigle and Kaufman (2006), and my study used topographic maps and aerial photographs
to derive modern ELA values. The location of the ‘modern’ ice front is a cumulative
product of the previous decade’s (post-LIA) climate, whereas the mass-balance studies
25
reflect the ELA of the limited study period. Therefore the higher ∆ELA values of Calkin
et al. (1985) may be attributed to the high ELA values of the 20th century.
The estimated ∆ELA in the Cascade Lake study area (23 ± 13 m) can be used to
estimate temperature lowering during the LIA. Using a typical environmental lapse rate
of 6°/km (Porter, 2001) and assuming no change in precipitation, results in a cooling of
0.13 ± 0.07 °C during the LIA. This lapse rate is only slightly lower than the regional
environmental lapse rate (7°/km) calculated from the summer temperature data in this
study. This amount of cooling is significantly less than the LIA cooling of 2-3°C
reported for the Brooks Range (Calkin et al., 1985), ≤ 1 °C temperature depression for the
Wrangell Saint Elias Mountains (Davi et al., 2003), ~1.7 °C cooling in the Alaska Range
(Hu et al., 2001), and the average 0.5-1.5 °C reported for the Northern Hemisphere
(Grove, 2003; Maasch et al., 2005). The significantly cooler temperatures of Calkin et al.
(1985) derive from their large ∆ELA values. The calculation of temperature cooling
from lapse rates and ∆ELA values is dependent on the assumption that no change in
moisture affected glacial expansion. However, if the northeastern Ahklun Mountains
were drier during the LIA, this would have limited glacier expansion despite cooler
temperatures. This scenario of a cold and dry LIA maximum in the study area may
explain the discrepancy in LIA temperature cooling calculated from lapse rate values.
26
CHAPTER 4: CASCADE LAKE SEDIMENTS
Background
Cascade Lake was targeted for a lake sediment study because it has a high
sediment load and is hydraulically connected to up-valley glaciers that show evidence of
late Holocene fluctuations. Additionally, Waskey Lake (Levy et al., 2004) to the south,
and Arolik Lake (Kaufman et al., 2003) to the west have been used successfully to
reconstruct paleoclimates in the region. Cascade Lake sediments were analyzed for a
variety of proxy indicators including their geochronology, magnetic susceptibility (MS),
organic-matter content (OM), biogenic silica (BSi), and lamination stratigraphy.
In lake sediments, MS is affected by organic content, density, grain size and
mineralogy (Dearing, 1999), and it is not always possible to discern which property is
forcing the MS signal. For this reason, MS is used to identify shifts in sedimentation
patterns that can be correlated between cores and in some cases attributed to climate
variability when used in combination with other proxy records. Variations in OM reflect
a number of factors, including lake productivity, minerogenic flux, the input of terrestrial
vegetation to the lake and diagenetic preservation. Because climate fluctuations can
directly affect each of these factors, OM is a valuable indicator of climate (e.g., Willemse
and Tornqvist, 1999; Kaufman et al., 2003; Bakke et al., 2005). BSi reflects the
abundance of diatoms in lacustrine sediments. Diatoms are single-celled algae that
typically dominate internal productivity in a lake and are sensitive to environmental
changes, making them good climate proxies (e.g., Willemse and Tornqvist, 1999;
Anderson, 2000). The deposition of laminations in a lacustrine setting can be driven by a
27
variety of hydrologic and meteorological controls (Leeman and Niessan, 1994;
Lamoureux, 1999). This relationship between climate and lamination thickness has been
used to produce quantitative and qualitative estimates of climate variability on annual to
seasonal scales (Hughen et al., 2000; Moore et al., 2001; Loso et al., 2006). Each of
these parameters were analyzed in Cascade Lake and interpreted for patterns that may be
indicative of climate variability.
Methods
Bathymetry, Acoustic Profiling, and Core-Site Selection. Core site selection was based on
bathymetric data and acoustic profiles. Bathymetric data were acquired with a Lowrance
sonar unit and subsequently plotted in GIS with the shoreline based on 1973 maps. The
sub-bottom stratigraphy was surveyed using an EdgeTech GeoStar acoustic profiler
operating between 4 and 24 kHz. This profiling identified regions of continuous and
condensed stratigraphy (avoiding slumps, methane, etc.). Three sites (distal-shallow,
proximal-shallow, and proximal-deep) were targeted for percussion coring (Figure 13).
Cores up to 6 m in length (7.5 cm diameter), were successfully recovered from these sites
(Table 6). Coring was completed from an anchored floating platform and was continued
until the core tubes were filled or an impenetrable layer was encountered. Because the
uppermost sediments in the long cores are often disturbed, surface cores were also
recovered from each site. These preserve the sediment-water interface and allow for an
evaluation of the stratigraphic completeness in the long cores. Surface cores were either
sub-sampled in the field into 0.5-cm increments or dewatered and packed. All cores were
shipped to Northern Arizona University and placed in cold storage.
28
Lithostratigraphy. Cores were split and photographed at ~10-cm intervals to document
sediment color prior to oxidation. General sediment descriptions (color, layers, bedding
thickness, texture, etc.) were logged. Unique features, such as tephra beds (described
below), were sub-sampled and viewed under a binocular microscope.
Additionally, a test pit was dug on the northern shoreline of Cascade Lake
(Figures 13). The stratigraphy of the test pit was described and sub-samples were taken
for radiocarbon analysis and visual comparison to the lake core stratigraphy.
Geochronology. A variety of methods were employed to determine the ages of the
lacustrine sediments, each with well-established practices. These included: radiometric
210Pb and 137Cs dating, tephrochronology, and radiocarbon dating. 210Pb analysis was
completed on surface core 02A at the University of Illinois by alpha counting and on
Core 02 at University of Southern California by gamma spectrometry in conjunction with
137Cs analysis. Contiguous 0.5-cm-thick (~1 cm3) samples were submitted for analysis.
Both 210Pb profiles were modeled using a constant rate of supply (CRS) model because
this model copes with variable sedimentation rates, which probably fluctuated
significantly (Appleby, 2004).
Tephras were located by their coarser grain size and their volcanic glass shards
and pumice. MS peaks were also targeted due to the association of volcanic glass and
magnetite. Each tephra was described under a binocular microscope to compare with
other tephras known from the region. I analyzed the major-element geochemistry of the
most prominent (visible and thickest) tephra in the Cascade Lake cores to correlate it with
29
tephra of a known age. The tephra was sampled from the center of the unit in Core 02 at
a depth of 2.85 m. It was cleaned with 2% NaOCl and dried. Glass grains were hand
picked under a binocular scope, epoxied onto a glass slide, polished with 1000 µm grit,
then carbon coated for analysis on a Cameca MBX electron microprobe at Northern
Arizona University. A single point on each of 31 glass grains was analyzed along with a
standard of basaltic glass (USM 11349811) used to calibrate the microprobe prior to use.
Total oxide percents of 10 elements were normalized to 100%. Similarity coefficients
(Borchardt et al., 1972) were determined by comparison with published values for the
3500 cal yr BP Aniakchak tephra (Riehle et al., 1987; Béget et al., 1992; Levy, 2002).
Lake sediment samples were sieved every ~5 cm down the core profile in search
of macrofossils, but no sample contained sufficient material for radiocarbon analysis. A
single radiocarbon age was obtained from a buried soil in the test pit on the north
shoreline of Cascade Lake. A sub-sample of the buried soil was wet-sieved and
handpicked for vegetation macrofossils. Clean and dry macrofossils were analyzed by
AMS at the University of California, Irvine. The radiocarbon age was calibrated using
CALIB v. 5.0 (Stuvier et al., 1998) and the reported age is the median probability age of
the 1σ calibrated age range.
Magnetic Susceptibility. MS was measured on the split core face using a Bartington MS2
meter, a hand-held MS2E1 detector, and Multisus computer software to correct for
measurement drift. Refrigerated cores were brought to room temperature before analysis.
Measurements were taken every 0.5 cm with drift correction every five readings. MS
values are reported in Systéme International (SI) units and were taken at a 0.1 sensitivity
30
range setting. Sediments with anomalously high MS values were examined under a
binocular microscope to confirm the presence of tephra shards, or unique mineralogy.
Organic-Matter Content. OM and bulk density were measured on two of the three cores.
Contiguous 1 cm3 samples were taken every 1 cm down to a depth of 198 cm in Core 01,
and over the entire length of Core 02. In a few places sample resolution was reduced
over coarse-grained, dense tephra units. Samples were processed using the procedure
described by Dean (1974). They were dried at 90°C for 12 hr, and then weighed to
measure dry bulk density. OM was approximated by LOI after combusting samples at
500°C for 5 hr. Because inorganic carbon content is negligible in the sediments, it was
not measured.
Lamination Stratigraphy. Although laminations were faintly visible on much of the
surface of the cores they were too diffuse for accurate measurement of thickness based on
visual inspection. Therefore, the core face was sub-sampled to a depth of 140 cm and
imaged by X-radiography. Slabs were 15 x 0.4 cm and were taken from the core axis
with ~1 cm of overlap between the slabs. These uniform-thickness slabs were X-rayed at
the NAU health center on detail film (3/20 sec, 50 mA, 50 kVp). Laminations were
visible on the X-rays as discrete zones of millimeter-scale, alternating high and low
density. Images were digitally spliced and lamination couplets were counted (four times)
and the thickness of both the high- and low-density layers and the total couplet was
measured (Figure 14).
31
Biogenic Silica. Biogenic silica (BSi) analysis was performed using the methods
described by Mortlock and Froelich (1989). Contiguous 2-mm-thick samples were taken
from Core 02 between 0 and 0.15 m, and from 0.15 to 1.50 m at a 5 mm sample interval.
The rest of the core was sampled in 5-mm-thick sections at 5-cm intervals. Hydrogen
peroxide (H2O2) and hydrochloric acid (HCl) were added to ~100 mg of dried and
powdered sample to remove organic material and carbonate. BSi was then extracted with
10% sodium carbonate solution (Na2CO3). The supernate containing dissolved silica was
then treated for molybdate-blue spectrophotometry through the reduction of the
silicomolybdate complex. BSi concentration was determined with a HACH DR/2000
spectrophotometer set to 812 µm wavelength and zeroed against a blank standard of
Na2CO3. Absorbance values were converted to BSi content using a standard curve
created by NAU lab technician Caleb Schiff (Fall 2005).
Samples were run in groups of 40 and experienced some batch effects
predominantly caused by slight variations in the mixing of the siliciomolybdate reduction
solutions. In order to adjust for these batch effects, three samples from the initial batch
were treated as standards and processed with every subsequent batch. Batches were then
calibrated to the original batch by adjusting BSi values by the average difference of the
three sample standards. Although the reported BSi content depends on these adjustments,
the relative difference among sampled is reproducible to ± 0.3% (1σ) based on multiple
runs of the standard samples.
32
Results and Discussion Bathymetry, Acoustic Profiling, and Core Sites. Cascade Lake contains a deep trough
(>50 m) running from the main inflow on the south end of the lake to the outlet on the
north. Broad, sloping deltas extend from the main and eastern inflows (Figure 13).
Acoustic profiling revealed generally continuous stratigraphy with localized zones of gas
disruption (Figure 15). The recovered cores did not penetrate the full stratigraphic
sequence visible in the acoustic profiles, with about 3 m of sediment remaining below the
base Core 01; however, there was little acoustic stratigraphy visible beneath the base of
Core 02. This could be because the stratigraphic sequence of this core is more complete
or because the acoustic signal was degraded.
Lithostratigraphy. Core 01, from the north distal core site, consists of five main
stratigraphic units with gradational contacts (Figure 16). (1) The basal unit (4.7 to 4.2 m)
is composed of sand interbedded with gray (5B 5/1) mud. The fine-grained layers are
separated by sharp contacts from the sandy mud. Unit 2 (4.2 to 3.6 m) is well-laminated
with laminations apparently defined by grain size (silt rich vs. clay rich). The
laminations range between ~0.25 and 1.0 cm thick. Unit 3 (3.6 to 2.8 m) consists of
brown (2.5Y 5/4, 106 5/6) clay-rich mud. Unit 4 (2.8 to 1.1 m) is blue-gray clay-rich mud
with three coarse-grained tephras layers. Unit 5 (1.1 to 0 m) is light-brown (2.5Y 5/4, 5Y
5/1) mud with very faint laminations in the uppermost ~30 cm.
The lithostratigraphy of Core 02 is very similar to Core 01 (Figure 16). It
comprises five fine-grained units defined by color differences and separated by
gradational contacts. (1) The basal unit (4.4 to 3.8 m) consists of sand interbedded with
33
gray, clay-rich mud. The sand layers are better defined than those of Core 01. Unit 2
(3.8 to 3.7 m) contains well-defined laminations > 1 cm thick, separated by sharp
contacts. Unit 3 (3.7 to 3.1 m) is composed of clay-rich, light-yellow-brown (10YR 5/6)
sediment with coarser-grained (not tephra) beds averaging 2 cm thick. Unit 4 (3.1 to 1.6
m) consists of blue and olive gray, fine-grained sediment (5B 5/1, 5Y 5/2) with five
coarse-grained tephras. Unit 5 (1.6 to 0 m) is olive to yellow-brown (2.5Y 4/4, 10YR
5/6, 2.5Y 6/2) mud with faint laminations throughout. These laminations are a centimeter
to millimeter thick.
Core 03 (deep, main basin core) is markedly different than Cores 01 and 02
(Figure 16). It contains well-defined, thick (0.5-3.0 cm) beds of alternating silt- to clay-
rich mud and disturbed and deformed sediments that interrupt the bedded stratigraphy.
Unit 1 (6.0 to 4.4 m) is composed of deformed sediments, coarse-grained beds and > 1-
cm-thick laminations. Unit 2 (4.4 to 3.2 m) consists of coarse-grained layers interbedded
with massive mud. Unit 3 (3.2 to 0.9 m) is predominantly deformed sediments with some
discrete units of mud. Unit 4 (0.9 to 0 m) is composed of well-defined laminations that
are discontinuous and separated by deformed sediments. Overall, Core 03 contains four
zones of well-formed, grain-size-defined laminations. No tephras were observed in the
Core 03.
The tops of Cores 01 and 02 were correlated with their sister surface cores using a
distinct fine-grained marker bed that grades upward from dark brown to red-brown. In
Core 01 this marker bed is at 8.0-8.5 cm depth, and at 8.5-9.0 cm in the surface core. In
Core 02 the bed is thicker and at a depth of 13.5-15.0 cm compared to 14.0-15.0 cm in
the surface core (Figure 17). Using these correlations, ~0.5 cm is missing from the
34
surface of each of the long cores and core-tube depths were adjusted accordingly so that
all reported depths are relative to the lake floor.
The lithostratigraphy of Cores 01 and 02 can be roughly correlated (Figure 16).
Both contain a thick (~15 cm) tephra (at 2.12 m in Core 01 and 2.85 m in Core 02), and
the basal sediments are similar in both cores. Laminations are better preserved in Core
02, though they are subtly visible on the surface of Core 01.
Because discrete horizontal beds are preserved in sections of Core 03, the
disturbance and deformation probably was not caused by the coring process. Instead, the
sediment deformation is likely caused by underflow currents or slumps within the deep
trough where the core was recovered. The thick tephra unit visible in Cores 01 and 02 at
2.12 and 2.85 m, respectively, was not found in Core 03. This, in combination with the
thick laminations of Core 03 (> 1 cm), indicates that sedimentation rate at the Core 03
site is significantly higher than that of Cores 01 and 02. For these reasons little further
work was completed on Core 03.
I focused most of my work on Core 02 because the laminations were better
preserved than in Core 01, likely due to its closer proximity to the main inflow, and it has
a longer sedimentary sequence and is less disturbed than that of Core 03. The proximal-
shallow water location of Core 02 also might be more sensitive to climate variability due
to the greater influence of the main inflow compared to the distal site of Core 01.
The test pit dug on the northern shoreline of Cascade Lake was 25 cm deep and
was subdivided into five units (Figure 18). (1) The base of the soil pit ended in a poorly
sorted, sub-angular to sub-rounded cobble gravel. This is overlain by Unit 2 (24.5-10.0
cm) a beige-orange to brown-orange tephra that grades from coarse to fine sand and
35
contains pumice shards visible with a hand lens. Unit (3) (10.0 to 4.3 cm) is a buried A-
horizon that is dark-brown to black. Unit 4 (4.3 to 2.6 cm) contains a discontinuous gray
tephra, with pumice lapilli, forming discrete lenses and mixed with organic material from
the buried soil below. Unit 5 (0-2.6 cm) is organic matter of living tundra.
210Pb and 137Cs. The two 210Pb profiles analyzed from cores 02 and 02A are quite similar
(Figure 19, Table A-1). In the core 02A, excess 210Pb activity levels off at 1.0 dpm/g at
~13 cm depth, resulting in an age of 1856 AD at 12.25 cm (CRS model) and an average
sediment flux of 0.08 g/cm2/yr. The profile indicates a shift in sediment accumulation
rate at ~8.25 cm (1948 AD). Below this point the sedimentation rate averages 0.48
mm/yr, compared with 1.76 mm/yr above it. This sedimentation-rate shift is only
constrained by four points at the base of the profile. The 210Pb profile from Core 02
results in an age of 1851 AD at a depth of 12.5 cm (Figure 19) similar to Core 02A. The
210Pb activity of Core 02 exhibits a shift in sedimentation rate from 0.40 mm/yr below
8.25 cm to 1.48 mm/yr above this point. This shift occurred a 1952 AD according to the
210Pb age model from Core 02, but is constrained by only two points. The 210Pb profile of
Core 02 was supplemented with 137Cs analyses from the same sediments. 137Cs peaked
with atmospheric nuclear testing in 1963, and therefore can be used as an independent
measure of sediment age. The activity of 137Cs shows a prominent peak from 8-6 cm
with a maximum at 6.25 cm (Figure 19). According to the 210Pb profile, a depth of 6.25
corresponds to an age of 1971, not 1963.
The two 210Pb models for core 02A and Core 02 differ slightly, especially in the
youngest sediments, but are similar in sediments older than the 1800’s. Both cores
36
exhibit a major shift in sedimentation rate at 1950 AD. Because the 210Pb profile produce
by gamma spectrometry was analyzed on sediments directly from Core 02, which was
also analyzed for other parameters, I focused on this profile and incorporated it into the
age model. The offset between the 210Pb and 137Cs ages determined from the same
sediments might suggest an age error in the 210Pb ages. However, 137Cs was only
analyzed every 0.5 cm, and the age of 1963 lies between the two points that define the
137Cs peak in this profile. Assuming that the 137Cs peak occurs at ~6 cm, then the 210Pb
ages are within ~4 yr.
Radiocarbon. A single 14C age was obtained from the top 1 cm of the buried soil in the
test pit on the lakeshore. The sample consisted of a mixture of detrital wood and leaf
fragments (UCI AMS# 23856). The sample yielded a radiocarbon age of 435 ± 15 cal yr
BP, which calibrates to 505 ± 20 cal yr BP. This is a maximum age for the tephra that
overlies this buried soil.
Tephrochronology. Five prominent tephras were recognized in Core 02 at 3.41, 3.33,
3.24, 2.85, and 2.59 m (Table 7, Figure 16). All of these tephras were also recognized in
Core 01 (except the uppermost tephra). In addition, two tephras were described from the
test pit dug on the northern lakeshore (Table 7, Figure 13).
Two distinct populations were recognized in the geochemical data from the 2.85
m tephra in Core 02 (Figure 20, Table A-2), one with 74 ± 3% SiO2 (n = 16) and one with
59 ± 4% SiO2 (n = 15). This bimodal composition could derive from either zonation
within the magma chamber or from the mixing of multiple eruptions into this single
37
layer. Similarity coefficients between these average data and published geochemical data
for the Aniakchak tephra are as follows: 0.75 with Riehle et al. (1987, for the high SiO2
group); 0.76 with Béget et al. (1992, for the high SiO2 group); and 0.75 with Levy (2002,
both SiO2 groups). However, four of the individual point analyses showed a significant
(>0.95 similarity coefficient) correlation with the Riehle et al. (1987) samples.
The geochemistry of the 2.85 m tephra correlates poorly with the reference
samples of the 3500 cal yr BP Aniakchak tephra. On the other hand, the bimodal
composition of this tephra is indicative of the Aniakchak tephra. Of the four tephras
located in Ahklun Mountain lakes and analyzed for geochemistry by Levy (2002), only
the Aniakchak tephra exhibits this bimodal composition. The Aniakchak tephra is also
the thickest and coarsest-grained tephra in the region (Kaufman et al., 2003; Levy et al.,
2004), as is the 2.85 m tephra in Cascade Lake. Scanning electron microscope images of
the 2.85 m ash from Cascade Lake and a sample of known Aniakchak tephra also show
similar shard morphologies (Figure 21). Both samples have fibrous pumice, blocky
clasts, and bubble-walled platelets, all of which were also identified by Riehle et al.
(1987) in their distal Aniakchak samples. For these reasons, I believe that the 2.85 m
tephra in Core 02 is the Aniakchak tephra (= 3500 cal yr BP).
The stratigraphic position of the Aniakchak tephra helps to correlate other
tephras. Assuming the Aniakchak tephra is correctly identified in the Core 02, then the
base of the core is ~11,000 cal yr BP based on the extrapolation of sedimentation rates.
Waskey Lake cores contain four other tephras, only one of which is younger than 11,000
cal yr BP (Levy et al., 2004) and therefore young enough to be in the Cascade Lake
cores. This tephra, Tephra B, is a “pure, 1-cm-thick, dark tephra…dominantly fine sand
38
to coarse silt…the glass shards are platy, brown, phenocrystic and microlitic…bubble-
wall junctions are present”. In both Waskey Lake and Arolik Lake (Kaufman et al.,
2003), Tephra B dates to ~6100 cal yr BP, and is the next recognized tephra below the
Aniakchak tephra. In the Cascade Lake cores, the 3.24 m tephra is the next lower and a
likely candidate for Tephra B. However, the tephras at 3.33 and 3.41 m also fit the
description of Tephra B and are possible correlatives, although none were analyzed for
major-element geochemistry.
Overall, more Holocene tephras are observed in the Cascade Lake cores than in
any other lake core from the Ahklun Mountains; however, the record is complicated. The
505 cal yr BP gray tephra in the test pit could not be located in the cores despite an
intensive search. Additionally, the tephra at 2.59 m is located only in Core 02, and was
not found in the test pit or in Core 01. Of the other lakes cored in the Ahklun Mountains,
only Little Swift Lake (Axford and Kaufman, 2004) contains two tephras older than the
Aniakchak tephra and younger than 12,000 cal yr BP. Both Arolik Lake (Kaufman et al.,
2003) and Waskey Lake (Levy et al., 2004) contain a single tephra over this interval.
This is inconsistent with the three tephras in Cascade Lake below the Aniakchak tephra.
The inconsistencies among these tephra records are difficult to explain. The deposition
of tephra onto a snow-covered landscape and ice-covered lake may retard deposition of
the tephra in an individual lake. Additionally, spatial patchiness has been documented in
the physical parameters of lake sediments (e.g., water content, organic-matter, pigment,
and phosphorus; Downing and Rath, 1988). Similar spatial heterogeneity might apply to
the deposition of tephra in lake sediments. The presence of the 2.59 m tephra in Core 02
but not in Core 01 indicates that tephras are not deposited uniformly within large lakes.
39
Age Model. An age model was constructed for Core 02 based on the 210Pb CRS model
and the correlated ages of the Aniakchak tephra and Tephra B. Because Tephra B was
not definitively located in Core 02 and could be any of the three tephras between the 3.24
or 3.41 m, two age models were created to bracket the range of these possible
correlations (Figure 22). The age model was subdivided into three segments. The basal
section is an assumed linear sedimentation between Tephra B and the base of the
Aniakchak tephra. The average of the Tephra B alternatives results in a sedimentation
rate of 0.2 mm/yr. The deposition of the Aniakchak tephra between 2.85 and 2.70 m was
interpreted as being instantaneous. The next segment between the Aniakchak tephra and
the 210Pb age at 12 cm was assumed to be linear, with a sedimentation rate of 0.8 mm/yr.
For the upper segment, the CRS 210Pb model was approximated by a 3rd-order polynomial
fit with increasing sedimentation rates upwards (from 0.4 mm/yr at the bottom to 1.5
mm/yr at the top).
The pronounced shifts in sedimentation rates can be reconciled with other
evidence. The factor-of-four lower sedimentation rate below the Aniakchak tephra
agrees with the sedimentation-rate shift that Levy (2002) documented in Waskey Lake.
The striking increase in sedimentation rate at the top of the core (post-1950) corresponds
to the high (21 m/yr) retreat rate of Cascade Glacier from 1935-1957 and a period of
accelerated warming across the arctic (Overpeck et al., 1997).
If the low sedimentation rate between the Aniakchak tephra and Tephra B is
assumed to continue to the base of the core, then the basal sediments are approximately
11,000 cal yr BP using an average linear sedimentation rate of 0.2 mm/yr. This
continuation of a linear, low sedimentation rate below Tephra B is similar to that
40
determined by Levy et al. (2004) for Waskey Lake and the basal age corresponds with the
a northwestern Ahklun Mountain deglaciation age of ~12,000 cal yr BP (Manley et al.,
2001).
Magnetic Susceptibility. The average background MS value in all three cores is ~30 x 10-
6 SI (Figure 23, Table A-3). Both Core 01 and 02 show elevated MS levels in the basal 1
m and discrete zones of high MS (> 50 x 10-6 SI). In Core 01, the peaks in MS occur at:
3.29 m (75.5 x 10-6 SI), 2.73 m (175 x 10-6 SI), 2.27 m (58 x 10-6 SI), and 2.1 m (412 x
10-6 SI). The sediment at 2.1 and 2.73 m contains tephra. There was no visible change in
sediment composition at the 2.27 m MS peak, but the MS peak at 3.29 m contains >80%
diatoms (primarily Cyclotella). MS peaks in Core 02 occur at: 3.35 m (105 x 10-6 SI),
2.84 m (343 x 10-6 SI), and 2.59 m (114 x 10-6 SI). The peaks at 2.59 m and 2.84 m
correspond with tephras (see Tephrochronology). The sediments at 3.35 m (6250 cal yr
BP) are similar to those at 3.29 m in Core 01; both contain a high percentage of
Cyclotella diatoms. MS peaks in Core 03 are all less than 50 x 10-6 SI, and values vary
between 7 and 47 x 10-6 SI.
Most of the MS peaks in the Cascade Lake cores coincide with tephras. The
presence of multi-domain magnetite in coarse-grained tephras yields an elevated MS
value when compared to the ambient lake sediments. The only MS peaks that do not
coincide with tephra are the ~1-cm-thick layers containing >80% diatoms at 3.29 m in
Core 01 and 3.35 m in Core 02. The association of elevated MS values with high diatom
concentrations is unexpected because diatoms are pure opaline silica and should not have
high MS. Possibly, the algal bloom responsible for the high concentration of diatoms
41
was also associated with increased growth of magnetobacteria. Magnetobacteria have
been shown to respond directly to the same environmental factors that affect general lake
productivity (Kim et al., 2005). Additionally, the draw-down of oxygen by a large algae
bloom could produce anoxic conditions favorable for magnetobacteria growth (Kim et
al., 2005). The general rise in background MS at the base of Cores 01 and 02 is related to
the lithologic variability of sandy mud interbedded with fine-grained beds. This rise
could reflect the mafic-rich sandy beds or the increased density of the fine-grained beds.
Organic-Matter Content. OM in Core 01 ranges from 3.1 to 5.0%, whereas in Core 02 it
varies between 1.4 and 4.7% (Figure 24, Table A-4). Bulk density ranges between 0.7 to
2.2, and 0.5 to 2.1 g/cm3, respectively, in the two cores. The OM profile for Core 02 has
relatively low OM values at the base of the core that rise rapidly to a broad peak at ~3.55
m (7250 cal yr BP); then OM values gradually decrease to the top of the core. Bulk
density shows little overall change in this core. OM and bulk density generally decrease
towards the core top in Core 01. The highest OM values are at 1.54 m (4.9%) and 0.9 m
(4.8%). Each of these peaks is constrained by multiple samples.
Bulk density and OM in Core 02 are weakly correlated (r = -0.42; p = 1.3 x 10-18).
The overall low OM values reflect the pervasive minerogenic input from the glacial
inflows and the relatively low lake productivity and vegetation growth on the surrounding
landscape. The inverse relationship between OM and bulk density probably is due to
increased sediment accumulation (high bulk density) diluting the organic matter. The
decrease in OM in Core 02 at 3.55 m occurs over the prominent lithologic change from
gray mud to yellow-brown mud with well-developed, thick laminations. The lowest OM
42
values at the base of Core 02 and at 2.85 m correlate with the coarsest sediments related
to a tephra and sandy mud in these intervals.
Laminations. A total of 316 lamination couplets were counted between 0.05 (core top)
and 1.45 m. These couplets averaged 4.2 mm in thickness, with the dense layers
averaging 1.8 mm and the less-dense layers 2.4 mm thick (Figure 25). Based on
observations from the core face, the dense (light) layers are associated with an increased
proportion of fine-grained (clay) sediment relative to the less-dense (dark) layers. Total
couplet thickness is more strongly related to the thickness of the less-dense layers (r =
0.88) than to the dense layers (r = 0.73). The thickness of the dense and less-dense layers
show no relationship to one another (r = 0.32). Lamination thickness is highly variable
ranging from 1 to 15 mm.
The discernable seasonal signal in the sediment-trap material suggests that
Cascade Lake laminations are annual (varves). However, their relatively large average
thickness and the geochronology suggest that this is not the case. The high variability in
lamination thickness, with the lack of a systematic down-core pattern of fluctuation (i.e.,
no section contains laminations persistently above or below average thickness), indicate
that lamination thickness is not a reliable proxy for climate fluctuation in this lake.
Applying the age model to the lamination counts indicates that the average
lamination encompasses ~6 yr. This is on the scale of the sub-decadal climate-regime
shifts that affect southwestern Alaska (AL, ENSO; Papineau, 2001). Changes in North
Pacific climate regimes have been shown to influence fluvial dynamics and likely
lacustrine sedimentation. For example, in southeastern Alaska, Neal et al. (2002)
43
demonstrated that ‘warm’ PDO regimes correlated with higher-discharge winters (more
precipitation falling as rain as opposed to snow) and ‘cool’ PDO’s were related to higher
summer discharge (precipitation stored as snow and released during the peak of the melt
season). This type of multi-annual shift in runoff may drive shifts in sediment storage
and therefore the formation of laminations in Cascade Lake. Hodgkins et al. (2003)
showed that proglacial outwash plains can serve as either net sinks or net sources of
sediment with the shifts in sediment storage driven by the runoff regime. This variation
in sediment storage (potentially caused by North Pacific climate regime shifts) may
explain the variations in sediment flux responsible for lamination thicknesses.
The deposition and preservation of laminations can also be caused by processes
within the lake. The occurrence of large magnitude underflows or short-lived
environmental events (e.g., storms) can dominate the sediment record of some lakes and
result in the deposition of non-annual laminations (Lamoureux and Gilbert, 2004). These
processes are often associated with sedimentary structures that can be related to turbidity
flows or slumps, none of which were observed in Cascade Lake. Rhythmic laminations
were also recognized by Lemmen et al. (1988) in a glacial-fed lake on Baffin Island.
They interpreted the laminations as being caused by changes in the rate of sediment input
to the lake and/or changes within the circulation of the lake. They also attributed the lack
of varves to low total sediment input to the lake and the extremely fine nature of the
sediment entering the lake. Cascade Lake has relatively high levels of summer
suspended load, though concentrations were never measured, but the particle size is small
(avg. 4 µm) and may result in a long sediment residence time in the water column and the
non-deposition of varves.
44
Biogenic Silica. BSi in Core 02 averages 2% (i.e. 20 mg/g) and ranges from 0.6 to 6%,
and shows significant variability including three prominent peaks (Figure 26, Table A-5).
BSi values are the lowest at the base of the core and rise sharply at 3.35 m (6.0%; 6275
cal yr BP). Between 3.35 and 2.65 m, BSi values decrease again (3.0%) before rising to
form a broad peak at 2.65 m (5.4%; 3500 cal yr BP). BSi values return to an average
1.6% and then peak between 1.18 and 1.03 m at 2.5%. The average BSi value towards
the top of the core (1.03 to 0.0 m) is 1.9%.
The sharp peak in BSi at 3.35 m encompasses the ~1-cm-thick layer of diatoms.
The deposition of these diatoms was probably short-lived (perhaps within a single
season) and not caused by a long-term climate fluctuation. The BSi peak at 2.65 m
occurs above the thick Aniakchak tephra deposited at 3500 cal yr BP. The deposition of
tephra within a lake can encourage algal blooms, especially in environments where Si is a
limiting factor (Telford et al., 2004). Although BSi was increasing just prior to the
deposition of the 2.85 m tephra, the peak is significantly higher than the extrapolation of
this rise, and the BSi recovers to a lower level within 0.4 m of the tephra, suggesting that
this spike in BSi was caused by the tephra deposition.
45
CHAPTER 5: BIOGENIC SILICA AND CLIMATE INTERPRETATIONS
Diatoms are excellent climate indicators because they have a short life span and
respond directly and sensitively to physical, chemical and biological changes to their
environment. Because the translation of climate fluctuations to diatom productivity is
indirect, several assumptions must be made in order to use BSi as a proxy for climate,
among these is the assumption that BSi concentrations are directly related to the
productivity of diatoms.
The diversity and abundance of a diatom community in a lake is dependent on a
variety of factors, including: temperature, nutrient load, length of ice cover, turbidity,
thermal stratification, and duration of spring and fall turnover (Lotter et al., 1999;
Anderson, 2000; Smol and Cumming, 2000; Douglas et al., 2004). Each of these
environmental factors is directly linked to meteorological forcing (temperature, wind,
precipitation, etc.), and, in turn, to climate processes. The relationship between diatom
productivity and the physical/chemical characteristics of lakes is not always straight
forward. For example, Bradbury et al. (2002) described the relationship between spring
temperature and diatom concentration for Elk Lake, Minnesota. They discovered that
cold springs could result in either a small diatom population through late ice-out and
rapid over-turning or a large diatom population as a result of a slow transition to summer
and a long over-turning period.
Nutrient load is perhaps the most complex factor affecting diatom growth.
Silicon, nitrogen, and phosphorous are the dominant nutrients affecting the growth of
diatom communities and can be derived from a variety of sources. In arctic and alpine
46
environments, silicon is primarily derived from groundwater and the dissolution of
diatoms at the sediment/water interface, nitrogen (N) enters a lake from soils and
vegetation, and phosphorous is regenerated from the lake sediment during the winter and
transported to the photic zone during spring circulation, but is ultimately supplied from
bedrock and eolian dust (Bradbury et al., 2002). The establishment and concentration of
alder (Alnus) within the catchment following deglaciation, strongly controls N
concentrations in southern Alaska lakes (Hu et al., 2001; Engstrom and Fritz, 2006). The
development of soil and vegetation therefore is a primary influence on diatom growth.
The connection between lake productivity and climate is further complicated by
the methods used to determine lake productivity. The use of BSi as a measure of lake
productivity is dependent on the assumptions that: (1) The BSi preserved in the lake
sediment is from algal growth rather than phytoliths, (2) there is little recycling of silica
(dissolution of diatoms at the sediment/water interface), (3) there is minimal grazing of
diatoms by higher trophic levels, (4) all algal species contributing to primary productivity
are preserved in the lake sediment, and (5) there is no significant change in the size and
silica concentration of the diatom species (Bradbury et al., 2002; Colman and Bratton,
2003). The validity of these assumptions is difficult to assess from the sediment record.
In studies of lake productivity from Lake Baikal some have proven to be inadequate. For
example, Qui et al. (1993) and Swann and Mackay (2006) both measured low BSi
content in sediment with high diatom concentrations based on point counts and other
techniques. This discrepancy was attributed to dissolution of the diatoms at the
sediment/water interface during periods of low sedimentation.
47
Variations in BSi content in Cascade Lake are believed to reflect fluctuations in
lake productivity. The lake is ice-free for approximately five months of the year, and
likely has a significant nutrient load entering the lake, enabling a seasonal diatom bloom.
The rather high sediment load of the lake may limit the area of the photic zone available
for diatom growth, but surface area of the lake is large enough to support a significant
diatom population. Diatom dissolution is dependent on the pH, temperature, salinity and
ionic composition of the water and sediments (Ryves et al., 2001). Temperature was the
only one of these parameters measured in the Cascade Lake; however, it is unlikely that
extreme values of pH, salinity or ionic composition would be maintained in the lake for a
significant period of time, and diatom dissolution is considered to be in the lake. The
pronounced difference in BSi content in the trapped Cascade Lake sediments between
2004-2005 and 2005-2006 indicates that BSi is sensitive to interannual differences in
weather conditions and therefore to longer term climate relationships. Because the
laminations in Cascade Lake were determined to not be annual, I focused on fluctuations
in BSi as the primary indicator of paleoclimate variability.
Methods
In order to determine which climate parameter is most strongly correlated with
BSi in Cascade Lake sediments, I compared the BSi flux with meteorological records.
The 3rd-order polynomial fit to the CRS age model was used to infer the ages of the high-
resolution BSi samples analyzed from the top (0-14 cm) of Core 02 and to interpolate the
sedimentation rates of each 0.2-cm-thick BSi sample. The average temporal resolution of
the samples was determined to be 3 yr. BSi values were converted to BSi flux (BSi % x
bulk density x sedimentation rate) using bulk density values that were linearly
48
interpolated between the 1-cm-thick samples analyzed for OM. This conversion to flux
accounts for the effects of variable sediment dilution. These flux values were then
correlated with meteorological parameters (e.g., mean annual temperature, June-July-
August average temperature) from the instrumental records at Dillingham (NCDC, 2005;
WRCC, 2006), discharge data from Nuyakuk River (70 km to the southeast; NWIS,
2006), and climate indexes from the North Pacific Ocean (e.g., North Pacific Index,
Aleutian Low Pacific Index; Ebbesmeyer et al., 1995; Mantua et al., 1997; Beamish et al.,
1997; NCAR, 2006). Correlations were made using both unsmoothed data and five-year
running averages of the climate data. The five-year running average was to account for
possible error in the 210Pb age model, smoothing in the sedimentary record (bioturbation),
and the prolonged influence of sub-decadal climate variability leading to ‘memory
effects’ (e.g., a rise in the groundwater table that sensitizes a landscape to runoff in the
years following). Temperature and precipitation data are discontinuous from the
Dillingham station with significant gaps in the 1970’s and 1930-1940’s. Temperature
averages were not used if more than three days were missing during a month and annual
precipitation values were acceptable only for years with no missing values (n-values in
Table 8).
Results
Correlations with BSi flux were made with 33 meteorological and North Pacific
climate parameters. Of these correlations, 11 were deemed of interpretive value (Table
8). BSi flux showed the highest correlation with Nuyakuk River discharge (r = 0.73, p =
2.2 x 10-6); the five-yr running average of mean May temperature (r = 0.70, p =9.16 x 10-
49
5); and the five-yr running average of the Aleutian Low Pacific Index (ALPI, r = 0.68, p =
1.37 x 10-7). BSi flux also correlated with the Pacific Decadal Oscillation Index (PDO, r
= 0.42, p =0.003), and the Pacific Northwest Index (PNI, r = 0.41, p = 0.003), but not
with the North Pacific Index (NPI, r = 0.17, p = 0.229).
Discussion
Controls on BSi in Cascade Lake. I attribute the relatively strong correlations between
BSi, Nuyakuk River discharge, and May temperatures to the hydroclimate and landscape
processes that control nutrient input to the lake. Because Cascade Lake does not support
an anadromous fish population, all nutrients must enter the lake from terrestrial sources
(fluvial, eolian, through-flow, groundwater, etc.). The majority of fluvial input to
Cascade Lake traverses the glacial outwash plain of Cascade Glacier with the exception
of the small western inflow (Figure 13). Glacial runoff is typically low in nutrient
content (Qui et al., 1993), especially during periods of glacier retreat when the rock-water
(ice) contact time is decreased (Lanfreniere and Sharp, 2005). Because of this, nutrient
input to Cascade Lake (and algal growth) is probably more strongly related to
precipitation than to glacier melt. Summer precipitation and spring snow-melt (winter
precipitation) infiltrate the soil and flow over the landscape carrying nutrients to the lake.
The gage on the Nuyakuk River measures the outflow of Nuyakuk and Tikchik Lakes (70
km southeast of Cascade Lake). Because less than 0.01% of the drainage basin is
occupied by modern glaciers, summer ablation has little influence on Nuyakuk River
discharge. There is no precipitation data from Cascade Lake, and Dillingham is more
coastal than Cascade Lake, therefore Nuyakuk River discharge is the best approximation
50
of precipitation at Cascade Lake. The influence of runoff on algal growth in a nutrient-
limited system has also been inferred for Grandfather Lake (Hu and Shemesh, 2003) and
Arolik Lake (Hu et al., 2003) in the Ahklun Mountains. While this terrestrial input
(related to precipitation) is likely the dominant source of Cascade Lake nutrients, some
nutrients are probably re-circulated within the lake and along the lake shoreline, driven
by wind mixing of sediments and erosion of the shoreline (e.g. Verschuren, 1999).
The relationship between May temperature and terrestrial runoff is more complex.
In 2004-2006, Cascade Lake experienced melt-out in the last week of May or the
beginning of June. Warmer May temperatures may result in earlier melt-out, rapid spring
overturning, and an extended spring growing season. Additionally, earlier melt-out may
result in increased overland-flow and near-surface through-flow entering the lake early in
the season. If the ground is still frozen and lacks vegetation cover, precipitation may take
a more direct, near-surface route to the lake.
In southwest Alaska, spring temperature and regional precipitation are directly
influenced by the AL, dominant feature of winter atmospheric and oceanic circulation in
the north Pacific Ocean. A strengthened AL tends to be associated with warmer and
wetter winters along the Bering Sea due to the increased tendency for storm systems to
pump warm air poleward (Rodionov et al., 2005). Additionally, of the climate modes
defined by Mock et al. (1998), the northeast Pacific negative type and the northcentral
Pacific negative type are the only two that result in positive winter temperature and
precipitation anomalies in southwestern Alaska, and these two modes feature a
strengthened AL.
51
The annual NPI is also a measure of AL strength, yet it shows no correlation with
the BSi content of Cascade Lake (r = 0.17, p = 0.229). Although the ALPI and NPI are
both measures of AL strength, they are negatively correlated (r = -0.48, p = 3.46 x 10-7)
because they relate to AL variations differently. The ALPI is a measure of the average
ocean area affected by low pressures (<100.5 kPa) from December-March, whereas the
annual NPI is the area-weighted sea-level pressure (Beamish, 1997; Tenberth and
Hurrell, 1994). Because the NPI is based on annual statistics, not only winter, the index is
also influenced by the ENSO cycles of the South Pacific and therefore is less directly
related to the climate regimes that affect algal growth at Cascade Lake. However, if the
NPI index is calculated over the same months as the ALPI (DJFM), then the two indices
are more strongly correlated (r = -0.73, p = 1.41 x 10-18) and the correlation between BSi
flux and the DJFM NPI five-year running average becomes significant (r = -0.42, p =
0.002). The ALPI shows moderately strong correlations with the Dillingham winter
precipitation and May temperatures (r = 0.33, p = 0.056, and r = 0.31, p = 0.011,
respectively).
Paleoclimate Interpretations. The AL is likely the dominant factor in the regional climate
that affects BSi at Cascade Lake, both in the last century and throughout the late
Holocene. If so, then BSi fluctuations in the sediment cores can be used to characterize
variations in the AL over the late Holocene with high BSi values being interpreted as
‘mild’ climates (moist and warm, with a strengthen AL). This interpretation of BSi flux
was completed at a sub-decadal resolution back to ~2000 cal yr BP based on the high-
resolution sampling interval (Figure 27). BSi flux shows three major fluctuations over
52
this period. Lake productivity was generally low from 2000 to 1450 cal yr BP (BSi flux
= 1.4 mg·cm-2·yr-1) and rose from 1450 to 1150 cal yr BP (2.6 mg·cm-2
·yr-1). BSi flux
decreased from 1150 cal yr BP to 1850 AD (1.9 mg·cm-2·yr-1) and subsequently increased
to its highest value in 1995 (4.4 mg·cm-2·yr-1). The pattern of these fluctuations (not the
absolute timing) corresponds to the general climate patterns of the late Holocene,
primarily the Medieval Warm Period (MWP), the Little Ice Age, and 20th century
warming. Additionally, the climate patterns of the LIA interpreted from BSi flux are
supported by the geomorphic evidence for glacial fluctuations. Low BSi flux in the
1800’s corresponds to ice advance in the study area, followed by subsequent ice retreat
and moraine stabilization in ~1860, and a return to increased BSi flux (Figure 27). The
high productivity during the mid-1990’s in Cascade Lake was unequalled at any time in
the last 2000 yr and may indicate that the modern strengthening of the AL is significantly
greater than during the past.
Few high-resolution records of climate continuously span the last few millennia in
Alaska. The most complete and comparable records are that of Hu et al. (2001) and Loso
et al. (2006) in southern Alaska (Figure 2), and Anderson et al. (2005) in the southern
Yukon territory. Hu et al.’s (2001) record of oxygen isotopes from Farewell Lake
(western Alaska Range) indicates a relatively warm climate from 850 to 1200 AD (1150-
750 cal yr BP). This agrees with the general varve thickening in Iceberg Lake (Loso et
al., 2006) from 1000 to 1250 AD (1000-800 cal yr BP) and the AL strengthening
interpreted at Jellybean Lake (Anderson et al., 2005) between 1200 and 900 cal yr BP.
This warm period is also indicated by the general lack of glacial advances across the
Northern Hemisphere (Porter, 1986). If the rise in BSi flux between 1450 and 1150 cal
53
yr BP in Cascade Lake is related to the mild climate of the MWP, then the timing of
MWP initiation is not synchronous and shows progressively later initiation eastward
across southern Alaska over a 450 yr period. Cascade Lake is ~900 km west of Iceberg
Lake and 600 km southwest of Farewell Lake (Figure 2). Warming at Cascade Lake was
initiated 200 yr prior to that at Farewell Lake, 400 yr prior to that of Iceberg Lake, and
250 yr prior to Jellybean Lake (Figure 28).
The time transgression of MWP initiation may be explained by three possible
causes, or some combination thereof. (1) Each of the climate records discussed above
derive from different proxy records (varve thickness in Iceberg Lake, isotopes in
Farewell Lake and Jellybean Lake, and BSi in Cascade Lake). While each of these
proxies record climate variations, they may each respond to different combinations of
temperature and precipitation. Depending on the climate thresholds necessary to produce
a measurable change, each proxy may record the initiation of different ‘stages’ of the
climate event. (2) The age model for Cascade Lake over this time period is based on a
linear extrapolation of sedimentation rates to the Aniakchak tephra deposited at 3500 cal
yr BP. Sedimentation rates likely vary with varying climate regimes, and the assumption
of a linear sedimentation rate over this extended period inserts some uncertainty in the
Cascade Lake age model. (3) The average position of the AL migrated westward during
the MWP.
Extending the climate interpretations beyond the last 2000 yr in Cascade Lake
suggests other general late Holocene climate trends (Figure 28). BSi flux prior to ~6000
cal yr BP is low and shows little variation. This may, reflect, the flux calculation using
the low sedimentation rate below Tephra B, or might indicate limited nutrient input to the
54
lake. Following deglaciation, the lack of vegetation (specifically the lack of nitrogen-
fixing alder) would limit nutrient levels in the drainage basin. OM values increase
between 11,000 and 7000 cal yr BP, indicating a milder climate following deglaciation
and the cool period between 13,000-12,000 cal yr BP that has been inferred from lowered
productivity in Arolik Lake during that time (Kaufman et al., 2001). Both Waskey Lake
(Levy et al., 2004) and Cascade Lake record the highest OM preservation between 6500
and 7500 cal yr BP. Waskey Lake also experience limited glacial input during this time,
making this period a likely candidate for the regional Holocene thermal maximum. This
is significantly younger than the average 10,000 cal yr BP age of the Holocene thermal
maximum across central and eastern Beringia (Kaufman et al., 2004).
Both Waskey Lake and Cascade Lake also record cooler climates between 3000
and 2500 cal yr BP, after the deposition of the Aniakchak tephra, marking the regional
initiation of Neoglaciation. This age is younger than initial Neoglacial advances
elsewhere in Alaska (4000-3500 cal yr BP in the Brooks Range, Ellis and Calkin, 1984;
4000-3000 cal yr BP on the Seward Peninsula, Calkin et al., 1998) and indicates a time-
transgressive climate shift potentially similar to that of the Medieval Warm Period.
Following the initiation of Neoglaciation, the glacial records at Waskey Lake and
Cascade Lake areas diverge, with multiple Neoglacial moraines in the headwaters of
Waskey Lake and only a single moraine formed late during the LIA around Cascade
Lake. As discussed previously, the cause of this discrepancy is uncertain, though it may
be related to variations in glacier flow mechanics caused by different subglacial bedrock
lithologies.
55
CHAPTER 6:
CONCLUSIONS
This study combines a moraine record of glacial fluctuations with an
interpretation of the sediments of Cascade Lake to determine the variability of Holocene
climate in the Ahklun Mountains. The glaciers of the Cascade Lake area reached their
maximum Neoglacial extent during the LIA. Lichenometric ages of five moraines
deposited by glaciers draining into Cascade Lake were derived using the Seward
Peninsula growth curve adapted from Solomina and Calkin (2003) with an increased
colonization time (~30 years). These ages place the stabilization of maximum LIA
moraines at ~1860 AD. Following moraine stabilization, Cascade Glacier retreated at a
variable rate, averaging 11 m/yr, to its current location 1.4 km up-valley.
During the LIA, cirque and valley glaciers of the Cascade Lake area experienced
an average ELA lowering of ~22 ± 12 m (n = 8) relative to 1957 AD. The ∆ELA values
of this study agree with other studies (Levy et al., 2004; Skiorski, 2004; Daigle and
Kaufman, 2006) that used similar methodologies to derive the ‘modern’ ELA. This
minimal ∆ELA in the Cascade Lake area probably reflects limited winter precipitation
during the LIA in southwestern Alaska.
The study of modern sedimentation patterns and temperatures of Cascade Lake
provides a context for interpreting late Holocene lake sediment. The contrast in sediment
accumulation and BSi flux between the two collection years (70% and 77% less,
respectively, in the cooler 2006 spring/summer) indicates the strong control of
spring/summer conditions on sediment and BSi accumulation in the lake. No
precipitation data are available for Cascade Lake or Dillingham, making it difficult to
56
determine whether this difference is related to temperature or to precipitation.
Precipitation has a more direct impact on discharge and terrestrial input of nutrients to the
lake, but both of these factors likely affect sediment and BSi accumulation to varying
degrees.
Annual fluctuations in the BSi, OM, and grain size analyzed in the trapped 2004-
2006 sediments indicate a seasonal signal is preserved in the sediments. Sediment with
low (high) BSi values correspond to elevated (decreased) bulk densities, and decreased
(increased) OM and grain size, and was probably deposited during the fall/winter
(spring/summer).
Of the three sediment cores recovered from Cascade Lake, I focused on Core 02
due to its longer record (compared to Core 03) and the occurrence of laminations. 210Pb,
137Cs, and the provisional identifications of the Aniakchak tephra and Tephra B provide
the age control for the core. The chronology below Tephra B is extrapolated using a
linear sedimentation rate. Overall, the chronologic control is limited and uncertain. MS
and OM values show relatively little variation down core compared to the flux of BSi,
and laminations are multi-annular (not varved). Therefore, I focused on BSi flux for the
purpose of paleoclimate reconstruction due to its apparent sensitivity to climate
variability and the ability to measure BSi at a high resolution (sub-decadal).
BSi flux (avg. 3 yr sampling resolution) was analyzed over the last 150 yr and
correlated with a variety of instrumental climate parameters. The climatological
parameters most strongly correlated with algal production in Cascade Lake are river
discharge at a nearby gaging station, spring temperature, and the strength of the AL, with
algal production favored during moist and warm conditions (strengthened AL). High BSi
57
flux indicates a mild climate from 1450-1150 cal yr BP (MWP?), a cool climate from
1150 cal yr BP to 1850 AD (LIA?), and a 20th century climate that was milder than that
of even the MWP.
If the interpretation of climate variation is extended to the full Holocene sequence
of lake sediments, other long-term trends can be inferred. The core did not extend into
ice-proximal sediments, so conditions immediately following deglaciation are not known.
Eventually the deglacial climate warmed and reached a period of maximum warmth
~7000 cal yr BP. Neoglacial cooling was initiated ~3000 cal yr BP, and glaciers reached
their maximum LIA extent during the late 19th century.
Interpretations of sediments from Cascade Lake and the glaciers in the
surrounding area indicate several periods of climate variability in the Ahklun Mountains
during the late Holocene. This variability is closely related to the dynamics of the North
Pacific Ocean and provides an important context for interpreting current and potential
future climate change. No other interval of the Cascade Lake sediment record is
comparable to the most recent period of warm and wet conditions, suggesting an
unprecedented strengthening of the AL during the 20th century.
58
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TABLE 1. INSTRUMENTS DEPLOYED AT CASCADE LAKE Instrument Location Latitude
Longitude (NAD 27)
Launch date
Recovery date
Comments
Mooring 1 Sediment trap Hobo Water
Temp Pro
Core site 01
~3 m water depth
Sediment trap
Hobo Water Temp Pro
Core site
01 ~23 m
water depth
60.14231 159.41209
6/28/04
8/12/05
8/10/05
7/19/06
Sediment accumulation =0.11 g/cm3/yr in
2004-2005
Mooring 2 Sediment trap Hobo Water
Temp Pro
Core site 03
~3 m water depth
Sediment trap Hobo Water
Temp Pro
Core site
03 ~51 m
water depth
60.13643 159.39917
6/28/04
8/12/05
8/11/05
7/19/06
Sediment accumulation =0.10 g/cm3/yr in
2004-2005 = 0.03 g/cm3/yr in
2005-2006
Air temperature Hobo Pro Western
lake shore (Fig. 3)
60.14082 159.41236
06/28/4
8/12/05
8/9/05
7/20/06
Temp. logger failed between 12/22/04 and 8/9/05
Hobo Pro Cascade Glacier ice proximal (Fig. 3)
60.08872 159.34043
8/13/05 7/20/06
69
TABLE 2. PHYSICAL AND BIOLOGICAL DATA FROM BOTTOM SEDIMENT TRAPS Unit Top
depth (cm)
Bottom depth (cm)
BSi
(%)
BSi flux
(mg·cm-2·yr-1)
Median grain size
(µm)
Organic matter (%)
Bulk density
(g/cm3) Mooring 1: 2004-2005 1 0.0 1.9 2.91 3.20 5.19 5.5 0.29 2 1.9 3.9 2.50 2.75 1.95 6.1 0.18 3 3.9 7.9 1.86 2.05 3.62 4.5 0.34 4 7.9 9.4 1.49 1.64 4.55 3.5 0.86 5 9.4 14.4 2.04 2.25 4.61 4.0 0.60 Mooring 2: 2004-2005 1 0.0 1.7 N.D. N.D. N.D. N.D. 0.40 2 1.7 8.9 N.D. N.D. N.D. N.D. 0.25 3 8.9 11.9 N.D. N.D. N.D. N.D. 0.44 4 11.9 12.9 N.D. N.D. N.D. N.D. 0.47 5 12.9 15.9 N.D. N.D. N.D. N.D. 0.31 6 15.9 17.7 N.D. N.D. N.D. N.D. 0.48 7 17.7 18.1 N.D. N.D. N.D. N.D. 0.66 Mooring 2: 2005-2006 1 0.0 1.6 1.27 0.38 N.D. N.D. 0.22 2 1.6 2.3 1.60 0.48 N.D. N.D. 0.37 3 2.3 2.5 1.59 0.48 N.D. N.D. 0.56 4 2.5 5.9 1.67 0.50 N.D. N.D. 0.11 5 5.9 8.1 1.30 0.39 N.D. N.D. 0.18 6 8.1 8.5 1.60 0.48 N.D. 0.66 Note: N.D. = no data collected on these samples.
70
TABLE 3. CLIMATE DATA FROM PREVIOUSLY PUBLISHED LICHEN
GROWTH-CURVE SITES OF SOLOMINA AND CALKIN (2003), CASCADE LAKE, AND OTHER LONG-TERM WEATHER STATIONS IN SOUTHWEST ALASKA
(DILLINGHAM AND BETHEL) Elevation
(m)
Mean annual temp.
(°C)
June, July, August (JJA) average
temp. (°C)
Mean annual precipitation
(mm)
Lichen growth-curve sites Brooks Range 350 -5.5 12.5 267 Seward Peninsula 4 -3.2 9.4 342 Wrangell-St. Elias 522 -5.7 13.3 201 Alaska Range 479 -2.9 12.6 239 Kenai (Homer) 27 3.0 11.1 518 Other climate stations
Bethel 36 1.4 12.0 430 Dillingham 20 1.0 11.7 647 Cascade Lake* 334 -4.0 8.7 Note: Climate data from the Spatial Climate Analysis Service (SCAS, 2005). *Based on daily averages at Cascade Lake from 2004-2006. Considered most similar to Seward Peninsula.
71
TABLE 4. SUMMARY OF LICHEN SIZES ON MORAINES AND OTHER DEPOSITS Glacier Latitude Longitude Deposit type Five largest
(mm) All measured
(mm) n Single
largest
Site #
(Fig. 8) (°N) (°W) Mean ±S.D. Mean ±S.D. (mm)
Lichen year (based on Seward Pen. growth-curve with ~30 yr
colonization time) Cascade 1 60.1018 159.3493 left lateral* 39 2 33 4 23 42 71
Cascade 2 60.1009 159.3485 left lateral* 29 1 25 4 20 30 115
Cascade 3 60.1020 159.3446 right lateral 32 4 26 6 13 39 129
Cascade 4 60.1019 159.3460 right lateral 34 4 29 4 18 40 140
E. Inflow 5 60.1314 159.3603 moraine crest† 36 3 33 3 14 39 152
E. Inflow 6 60.1339 159.3653 debris flow† 26 1 23 4 8 28 10
E. Inflow 7 60.1343 159.3734 debris flow† 27 4 23 4 19 32 106
E. Inflow 8 60.1344 159.3647 debris flow† 28 4 25 4 12 32 111
Hungry Joe 9 60.1066 159.3235 stable outwash 23 4 23 4 3 33 91
Divided 10 60.1189 159.3215 slump 26 3 23 4 6 29 102
Chikuminuk 11 60.1495 159.3127 river terrace 33 9 27 9 10 49 135
Cascade 12 60.0942 159.3414 stable outwash lichen/no
lichen contact
N.A. N.A. 6 1 4 6 49
Cascade 13 60.0956 159.3426 stable outwash 16 6 13 5 10 27 70
*Shown in Figure 10A. †Shown in Figure 10B.
72
TABLE 5. EQUILIBRIUM-LINE ALTITUDE (ELA) FOR GLACIERS IN THE CASCADE LAKE AREA Glacier Latitude Longitude 1957
length 1957 area
Head altitude
Toe altitude
Aspect 1957 ELA
Neoglacial ELA
ELA lowering
(see Fig. 8) (°N) (°W) (km) (km2) (m) (m) (m) (m) (m)
Cascade Glacier 60.0917 159.3380 3.8 3.76 1250 550 N 980 945 35
Divided Cirque 60.0951 159.3119 1.7 1.24 1190 730 NW 900 890 10
Hungry Joe Cirque 60.1066 159.3196 0.4 3.17 1190 980 SW 1045 1030 15
East Inflow Cirque 60.1311 159.3574 0.5 3.08 1070 820 NW 1085 1060 25
Chikuminuk Glacier 60.1300 159.2973 3.5 5.98 1280 610 NW 965 945 20
Cirque 1 60.0816 159.2999 0.6 1.69 1190 1010 NE 1045 1040 5
Cirque 2 60.0826 159.2843 0.5 2.78 1100 940 SE 1010 965 45
Cirque 3 60.1398 159.3378 1.0 4.08 1160 670 S 820 795 25
Average 980 960 23
Std. deviation 90 80 13
73
TABLE 6. LOCATION AND OTHER INFORMATION FOR LAKE CORES RECOVERED FROM CASCADE LAKE
Core Latitude (°N)
Longitude (°W)
Water depth (m)
Length (m)
Notes
04CA-01 60.1432 159.4167 24.2 4.68 04CA-01A 60.1432 159.4167 24 0.20 Sub-sampled in field
04CA-01B 60.1432 159.4167 24 0.46
05CA-01A 60.1435 159.4166 23 0.20 Sub-sampled in field
05CA-01B 60.1435 159.4166 23 0.46
05CA-01C 60.1435 159.4166 23 0.35
04CA-02 60.1327 159.3988 34.5 4.37 Some material lost from base during core recovery
04CA-02A 60.1327 159.3988 34 0.20 sub-sampled in field
04CA-02B 60.1327 159.3988 34 0.38
05CA-02D 60.1333 159.3995 34 0.20 Sub-sampled in field
05CA-02E 60.1333 159.3995 34 0.38
05CA-02F 60.1333 159.3995 34 0.31
04CA-03 60.1371 159.4039 52.4 6.03
04CA-03B 60.1371 159.4039 52 0.30
Note: Surface cores indicated by letter designations.
74
TABLE 7. TEPHRA DESCRIPTIONS FROM CORE 02 AND THE LAKESHORE
TEST PIT Basal depth (cm)
Thickness
(cm)
Glass type and color Bedding Contacts Notes
Core 02
2.59 1.1 Pink to cream bubble-walled shards
N.A. Sharp upper and lower
Not located in Core 01
2.85 15.0 Fine gray bubble-walled shards, coarse white pumice, significant microlitic glass shards
Reversely graded at base
Gradational upper and lower
Inferred to be the Aniakchak tephra (3500 cal yr BP)
3.24 3.0 Rounded with bubble wall junctions, clear to brown
N.A. Sharp upper and lower
Diatoms present
3.33 2.5 Brown glass shards N.A. Gradational upper, sharp lower
3.41 2.5 Brown glass shards N.A. Gradational upper, sharp lower
Test pit
4.30 2.0 Gray to white bubble wall shards
Preserved in pods
N.A. Overlies buried soil with a 14C age of 505 ± 20 cal yr BP
15.40 5.4 Orange/beige bubble wall shards with white pumice
Reversely graded at base
Gradational top, sharp bottom
Note: test pit illustrated in Figure 14.
75
TABLE 8. CORRELATIONS BETWEEN BSi FLUX AND METEOROLOGICAL AND PACIFIC OCEAN CLIMATE INDICES Correlated parameter* Unsmoothed 5-year running avg. Source† n-years R p-value n-years R p-values Annual precipitation 19 0.35 0.144 18 0.44 0.437 1
May precipitation 19 0.03 0.892 14 -0.10 0.733 1
Mean annual temp 32 0.14 0.447 26 0.11 0.576 1
Mean DJF temp 26 -0.04 0.828 21 -0.35 0.120 1
Mean JJA temp 21 0.26 0.260 17 0.36 0.159 1
Mean may temp 31 0.54 0.002 25 0.70 9.16 x 10-5 1
Nuyakuk River discharge 26 0.73 2.22 x 10-5 23 0.71 1.27 x 10-4 2
PDOI 51 0.31 0.028 49 0.42 0.003 3
NPI 52 0.04 0.760 50 0.17 0.229 4
DJFM NPI 52 -0.42 0.002 50 -0.42 0.002 4
ALPI 51 0.40 0.004 50 0.68 1.37 x 10-7 5
PNI 53 0.32 0.022 52 0.41 0.003 6 *DJF = December, January, February; JJA = June, July, August; PDOI = Pacific Decadal Oscillation Index; NPI = North Pacific Index; DJFM = December, January, February, March; ALPI = Aleutian Low Pacific Index; PNI = Pacific Northwest Index. †(1) WRCC, 2006 and NCDC, 2005; (2) NWIS, 2006; (3) Mantua et al., 1997; (4) NCAR, 2006; (5) Beamish et al., 1997; (6) Ebbesmeyer et al., 1995.
76
Figure 1: Periods of Holocene climate change. Columns denote the six periods of rapid climate change (polar cooling, tropical aridity)
proposed by Mayewski et al. (2004). Black boxes are centered over periods of climate cooling determined by various authors (left
column), and white boxes indicate mild, warm periods recognized by O’Brien et al. (1995).
77
Figure 2: Location of the Ahklun Mountains and features discussed in this thesis
(medium shading indicates elevations above 200 m). Lakes are indicated by numbers: (1)
= Cascade Lake; (2) = Waskey Lake; (3) = Arolik Lake; (4) = Little Swift Lake; (5) =
Farewell Lake; (6) = Iceberg Lake. Other locations indicated by letter designations: (AR)
= Alaska Range; (B) = Bethel; (BR) = Brooks Range; (BB) = Bristol Bay; (D) =
Dillingham; (KM) = Kenai Mountains; (KB) = Kuskokwim Bay; (NR) = Nuyakuk River;
(SP) = Seward Peninsula; (WSEM) = Wrangell-St. Elias Mountains.
78
Figure 3: Cascade Lake area. 1983 aerial photograph with arrows denoting inflows to
Cascade Lake; the western inflow is not glacier-fed. M1=Mooring 1, M2=Mooring 2,
AT=Air temperature station, GT=Glacier temperature station.
79
Figure 4: Instrumentation deployed at Cascade Lake. (A) Sediment trap and water-
temperature logger mounted on mooring line. (B) Installing air-temperature logger
housed in a solar-radiation shield in a stand of willow along the shoreline of Cascade
Lake.
80
Figure 5: Lithologic descriptions of the sediments collected in traps from Cascade Lake on August 10, 2005 and July 19, 2006.
Lithostratigraphy of bottom sediment traps from both mooring sites with physical and biological data (listed in Table 2).
81
Figure 6: Daily average water temperature in Cascade Lake 2004-2006. Bottom loggers
were positioned ~1 m above the lake bottom; top loggers were ~3 m below the water
surface.
82
Figure 7: Comparisons between average daily temperature at Cascade Lake, Dillingham, and Bethel, AK. (A) Differences between
sites. (B) Relationship between sites; note the slight seasonality effect shown by the increased offset from the 1:1 line for the coldest
winter temperatures.
83
Figure 8: Cascade Lake area showing glaciers referenced by informal names
(CA=Cascade Glacier, C1=Cirque 1, C2=Cirque 2, DIV=Divided Glacier, HJ=Hungry
Joe, CH=Chikuminuk Glacier, EI=East Inflow Cirque, C3=Cirque 3). Based on 1:63,360
Bethel 1979 topographic map with corrections for ice extent based on 1957 aerial
photographs.
84
Figure 9: Hypsometric curves used in the calculation of equilibrium-line altitudes (ELA)
for eight glaciers in the study area. An accumulation-area ratio (AAR) of 0.6 (60% of the
glacier in the accumulation zone, shown by the line at 40% of the glacier area) was used
for ELA calculations. ELA calculations (and especially ∆ELA) are insensitive to AAR
changes as shown by the relatively flat slopes between AAR values of 0.5-0.7 (shown by
the box).
85
Figure 10: Little Ice Age (LIA) moraines in the Cascade Lake drainage. Dotted lines highlight moraine crests and trim lines. (A) Left
and right lateral LIA moraines of Cascade Glacier. Deposits in the foreground are higher than those in the distance. (B) The East
Inflow cirque valley. The ridge crest closest to the cirque (dotted line) is the LIA moraine; the fan-shaped feature marked by ridges
below are likely debris-flow deposits. Numbers in white denote lichen stations. Photos taken in August 2005.
86
Figure 11: Growth curves used to derive the ages of measured lichens in the Cascade
Lake region (from Solomina and Calkin, 2003). The arrows show the average of the five
larges lichen diameters on moraines in the Cascade Lake area. Based on Seward
Peninsula curve, their ages average ~1860 AD assuming a ~30-yr colonization time.
87
Figure 12: Retreat rates for Cascade Glacier since the Little Ice Age based on aerial
photography (1984 and 1957) and lichen stations (1935 and 1860; see text).
88
Figure 13: Bathymetry of Cascade Lake with locations of coring sites. The main inflow
is in the south and the outflow is to the north.
89
Figure 14: Top 15 cm of Core 02 showing the relationship between the core surface
stratigraphy, X-radiograph stratigraphy, and the measurement of lamination thicknesses.
Lines connect points identified as being the same on the core face and X-radiograph
stratigraphy.
90
Figure 15: Acoustic profile image in the vicinity of core site 02. The white box denotes
the approximate location and length of Core 02. The acoustic signal is degraded at the
approximate base of Core 02, but stratigraphy deeper than this was observed in other
profiles and below Core 01.
91
Figure 16: Lithologic logs of three cores recovered from Cascade Lake.
92
Figure 17: Correlations between surface and long cores based on marker beds (dashed lines) in (A) Core 01 and (B) Core 02. Depths
for surface core refer to the sediment-water interface (lake bottom), and are within 0.5 cm of correlative depths in long cores.
93
Figure 18: Test-pit wall on the northern shore of Cascade Lake (Figure 13). * denotes
location of 14C sample dated at 505 cal yr BP.
94
Figure 19: 210Pb and 137Cs profiles from Core 02A and Core 02. The 3rd-order
polynomial was used to assign ages to the upper 12 cm of Core 02. Gray box shows the
intersection of 137Cs peak with ages derived from the 210Pb profile.
95
Figure 20: Geochemistry data from the 2.85 m ash in Core 02 and published data sets of
known Aniakchak tephras. Data listed in appendix A-2.
96
Figure 21: Scanning electron microscope images of tephra shards from (A) the 2.85 m tephra in Core 02 and (B) a known Aniakchak
tephra from Arolik Lake.
97
Figure 22: Age model for Core 02. The two models are based on the maximum depth
range of Tephra B candidates in Core 02 (see text for explanation).
98
Figure 23: Magnetic susceptibility (MS) profiles for Cores 01, 02 and 03 from Cascade
Lake. * indicates MS values that exceed 90 x 10-6 SI and dashed lines show correlations
between tephras.
99
Figure 24: Organic-matter (OM) content and bulk density profiles for the upper 2 m in
Core 01 and the entire Core 02.
100
Figure 25: Lamination thickness in the top 1.4 m of Core 02. Of the 316 lamination
couplets identified, the average thickness is 0.42 cm (bold line).
101
Figure 26: Biogenic silica profiles for Core 02. Inset shows detail of upper 14 cm, the
interval sampled at ~3 yr resolution and used to correlate with instrumental climate
records.
102
Figure 27: Biogenic silica flux from Core 02 over the last 2200 cal yr BP with periods of
recognized warm climate from Iceberg Lake (Loso et al., 2006), Farewell Lake (Hu et al.,
2001), and Jellybean Lake (Anderson et al., 2005).
103
Figure 28: Core 02 proxy data compared to climate events of the Yukon (Jellybean Lake; Anderson et al., 2005) and Waskey Lake
(Levy et al., 2004). OM = organic matter; BSi = biogenic silica. Gray bars indicate the Holocene Thermal Maximum (7000 cal yr
BP), Neoglaciation onset in the Ahklun Mountains (3000 cal yr BP) and Medieval Warm Period (~1000 cal yr BP).
104
TABLE A-1. 210Pb AND 137Cs DATA
Depth Unsupported 210Pb activity
137Cs activity Age of base
(cm) (dpm/g) (± s.d.) (dpm/g) (±s.d.) (years before 2004) Core 02A 0.5 9.16 0.53 5.48 1.5 3.80 0.37 10.95 2.5 4.29 0.26 17.58 3.5 3.88 0.23 27.31 4.5 1.88 0.17 35.90 5.5 1.31 0.19 43.22 6.5 0.45 0.16 47.38 7.5 0.46 0.16 51.13 8.5 0.51 0.15 55.88 9.5 1.30 0.24 70.45 10.5 0.57 0.17 88.45 11.5 0.48 0.14 113.58 12.5 0.22 0.14
N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. N.D. 148.35
Core 02 1.5 7.10 0.78 0.30 0.13 2.86 2.0 8.28 0.57 0.56 0.08 6.74 2.5 6.27 0.69 0.71 0.10 10.04 3.0 6.52 0.54 0.60 0.08 13.80 3.5 5.70 0.81 0.83 0.13 17.53 4.0 3.37 0.78 0.73 0.13 20.01 4.5 3.29 0.57 1.22 0.10 22.51 5.0 3.51 0.58 1.14 0.10 25.59 5.5 2.28 0.78 1.78 0.14 27.71 6.0 2.33 0.59 1.95 0.10 30.21 6.5 2.92 0.73 5.58 0.16 33.43 7.0 3.27 0.70 4.20 0.14 42.11 8.5 1.90 0.59 1.10 0.10 52.51 10.0 2.53 0.53 0.30 0.10 87.27 13.0 1.14 0.51 0.00 0.1 153.57
105
TABLE A-2. MAJOR ELEMENT GEOCHEMISTRY OF THE 2.85 m TEPHRA IN CORE 02
Point Na2O MgO Al2O3 SiO2 P2O5 K2O CaO TiO2 MnO FeO Total
1 2.397 0.489 15.370 72.958 0.069 2.651 1.706 0.439 0.152 2.392 98.623 2 2.228 0.488 15.429 72.122 0.057 2.415 1.767 0.439 0.094 2.385 97.424 3 3.302 2.605 16.468 59.778 0.784 1.409 5.906 1.306 0.186 6.800 98.545 4 1.290 0.489 15.514 72.879 0.096 2.225 1.797 0.394 0.097 2.357 97.137 5 2.708 3.220 16.466 57.271 0.807 1.131 7.146 1.443 0.194 8.197 98.583 6 1.525 1.323 16.211 64.919 0.538 1.449 3.839 1.093 0.090 3.760 94.748 7 0.997 0.464 15.582 74.064 0.092 1.718 1.820 0.445 0.105 2.434 97.722 8 1.645 2.225 16.145 62.412 0.577 1.508 5.334 1.253 0.142 6.174 97.420 9 1.529 2.784 16.519 58.862 0.754 1.191 6.361 1.409 0.178 7.299 96.888 10 1.089 0.439 15.242 72.922 0.103 1.523 1.767 0465 0.085 2.217 95.853 11 1.665 2.406 16.665 61.372 0.747 1.311 5.706 1.328 0.159 6.429 97.787 12 2.304 2.555 16.635 60.803 0.655 1.229 6.249 1.308 0.223 7.166 99.126 13 1.988 2.920 16.026 56.595 0.658 1.114 6.398 1333 0.213 8.197 95.443 14 1.007 0.479 15.537 73.401 0.137 1.533 1.854 0.435 0.121 2.417 96.923 15 2.141 2.793 16.298 58.920 0.644 1.119 6.287 1.323 0.159 7.402 97.085 16 0.737 0.363 12.259 56.077 0.057 1.075 1.545 0.379 0.090 2.120 74.702 17 0.743 0.406 12.763 61.909 0.050 0.928 1.591 0.395 0.039 2.240 81.064 18 1.851 1.711 16.973 64.354 0.435 1.444 4.587 1.091 0.156 4.729 97.331 19 2.188 0.425 15.680 74.205 0.028 1.957 1.692 0.484 0.110 2.428 99.196 20 1.334 0.428 15.369 73.521 0.147 1.788 1.713 0.470 0.141 2.304 97.214 21 3.169 3.365 16.294 57.525 0.603 0.875 7.251 1.281 0.152 8.807 99.322 22 2.668 2.849 16.595 59.517 0.726 1.024 6.049 1.358 0.199 7.033 98.017 23 3.193 2.899 16.655 58.888 0.758 0.890 6.702 1.433 0.199 7.032 98.650 24 1.587 0.839 15.478 71.489 0.034 1.629 2.695 0.570 0.077 3.061 97.459 25 1.256 0.491 15.828 73.831 0.034 1.636 1.990 0.445 0.114 2.366 97.991 26 2.564 2.577 16.963 60.082 0.625 1.012 5.928 1.528 0.117 6.587 98.042 27 1.349 0.413 15.524 73.514 0.000 1.683 1.880 0.460 0.124 2.556 97.504 28 1.426 0.463 14.704 70.288 0.050 1.432 1.763 0.434 0.119 2.213 92.892 29 1.086 0.446 14.908 71.677 0.069 1.160 1.851 0.439 0.094 2.287 94.017 30 2.796 3.113 16.287 57.436 .736 0.790 6.657 1.311 0.196 7.468 96.789 31 1.277 0.486 15.331 73.576 0080 1.390 1.840 0.429 0.085 2.404 96.898 std. 2.587 5.051 12.431 50.532 0.509 0.729 9.472 3.900 0.198 13.473 98.881
106
TABLE A-3. MAGNETIC SUSCEPTIBILITY DATA FOR CORE 02
0 20 40 60 80 100 120 140 160 180 200 0.0 27.6 33.8 27.0 27.0 23.7 27.0 24.6 28.1 27.0 23.5 0.5 27.9 27.8 30.9 26.0 28.0 24.9 26.0 22.8 26.4 28.0 25.3 1.0 26.9 27.2 28.4 24.9 27.8 26.0 28.0 22.2 26.7 26.0 31.2 1.5 26.0 33.4 27.5 31.7 27.7 24.0 26.0 20.2 32.0 26.1 26.9 2.0 23.0 31.5 30.6 26.5 26.5 26.0 25.0 15.3 29.0 24.3 21.7 2.5 23.0 29.7 29.7 24.3 29.4 31.0 26.0 16.4 28.0 23.5 22.5 3.0 27.0 33.9 28.9 26.2 28.2 29.0 25.0 23.5 28.0 24.6 24.4 3.5 19.0 34.2 30.0 28.0 28.7 29.0 24.2 23.6 27.0 26.8 26.2 4.0 22.0 34.3 29.0 23.0 26.4 28.0 24.4 23.6 25.2 26.0 21.8 4.5 26.0 30.5 29.0 25.0 25.0 28.0 24.6 22.7 17.6 26.0 20.6 5.0 35.0 26.7 36.0 28.0 25.7 25.0 23.7 23.8 25.5 26.0 21.5 5.5 29.0 27.9 27.0 30.0 27.3 24.0 23.9 23.9 28.7 24.0 24.3 6.0 28.8 24.0 25.0 28.0 29.0 28.2 26.7 20.3 26.4 25.0 24.2 6.5 32.0 25.0 26.0 25.0 28.0 24.3 26.6 20.5 24.5 24.6 28.7 7.0 31.3 25.0 30.0 24.0 28.0 24.5 27.4 23.6 27.6 24.5 25.6 7.5 31.5 31.0 26.0 23.0 27.0 23.7 27.3 21.7 31.7 23.4 24.5 8.0 33.8 32.0 23.5 23.0 28.0 25.8 28.2 22.9 29.8 21.3 26.4 8.5 29.1 33.0 26.6 23.2 28.0 25.1 30.0 28.1 26.9 22.1 31.2 9.0 27.3 30.0 33.7 24.3 26.0 29.3 29.0 25.1 30.0 25.0 29.7 9.5 29.5 24.0 25.7 25.5 25.0 31.5 27.0 18.2 29.0 25.0 30.6 10.0 27.6 24.0 26.8 25.7 23.0 25.7 28.0 18.5 29.0 25.0 27.4 10.5 27.8 28.0 27.9 27.8 26.0 26.8 26.0 18.9 29.0 26.0 23.3 11.0 28.1 28.2 28.9 26.0 26.0 26.1 27.0 19.8 31.0 23.0 25.1 11.5 26.3 29.3 26.7 26.0 24.0 28.3 26.0 16.0 32.0 24.9 27.6 12.0 24.5 29.5 28.5 23.0 25.0 25.5 25.0 24.0 31.0 24.7 27.5 12.5 25.7 29.7 29.4 23.0 23.0 25.6 27.0 27.0 27.0 26.5 25.4 13.0 20.9 31.9 27.2 22.0 26.0 26.8 24.0 26.0 26.0 23.4 25.2 13.5 20.2 32.0 27.8 27.0 27.0 24.0 23.0 27.0 27.0 23.2 26.1 14.0 22.4 29.0 29.6 23.0 26.0 24.0 23.0 24.3 30.8 21.8 26.8 14.5 24.6 29.0 31.5 24.0 27.0 24.0 26.0 25.6 27.6 21.6 26.7 15.0 28.7 30.0 29.3 26.0 27.0 25.0 25.0 27.0 26.5 22.5 31.5 15.5 28.9 27.0 27.1 25.0 30.0 25.0 24.0 27.3 26.3 23.3 33.3 16.0 29.8 24.1 25.8 30.1 29.2 26.0 24.4 25.7 29.1 24.1 31.2 16.5 26.7 28.3 29.6 25.3 26.3 23.0 24.5 24.5 27.0 24.5 27.0 17.0 25.5 30.5 27.5 24.5 27.5 25.0 25.6 25.0 29.0 23.4 26.0 17.5 25.4 29.6 27.3 25.6 27.6 27.0 21.8 28.5 27.0 24.3 25.0 18.0 29.2 29.8 24.2 26.8 26.8 20.0 22.9 30.0 28.0 24.2 27.0 18.5 27.1 31.4 25.0 25.0 24.3 29.0 22.2 26.4 26.0 25.1 27.0 19.0 29.3 32.5 28.0 28.0 23.4 28.0 21.3 26.3 27.0 27.8 27.0 19.5 27.4 34.6 27.0 29.0 25.6 28.0 24.5 26.6 28.0 24.7 29.0
107
220 240 260 280 300 320 340 360 380 400 420 0.0 29.0 30.4 55.5 289.5 22.0 21.2 17.0 15.6 17.0 27.2 23.0 0.5 25.0 28.3 58.4 227.4 21.0 18.4 16.0 16.5 18.0 23.3 23.0 1.0 26.0 27.1 63.2 204.2 20.0 15.5 17.0 15.4 17.0 15.5 31.0 1.5 28.9 27.8 65.7 241.9 19.0 14.7 17.0 15.3 17.0 15.7 30.0 2.0 29.7 25.7 73.6 220.8 17.0 14.9 16.0 14.1 15.0 21.8 26.0 2.5 25.5 25.5 100.4 274.7 16.0 16.2 17.0 16.0 16.1 26.2 26.7 3.0 22.3 27.4 136.3 299.6 16.0 19.4 16.0 17.0 17.3 25.4 25.8 3.5 21.2 31.2 115.1 332.5 17.0 21.5 16.0 16.0 17.4 23.5 26.8 4.0 22.2 29.8 119.8 312.5 18.0 21.7 16.0 15.0 18.6 23.7 30.9 4.5 24.4 25.7 106.7 314.3 17.0 20.9 17.0 14.0 19.8 19.8 30.9 5.0 25.5 23.5 141.5 287.2 16.0 17.3 18.0 14.1 21.9 19.1 30.2 5.5 24.7 23.4 156.3 143.2 16.0 16.6 17.0 14.3 21.7 22.3 32.3 6.0 23.8 22.2 161.2 120.1 17.0 17.0 16.0 16.5 20.5 16.5 34.5 6.5 23.0 25.7 128.1 131.0 14.0 18.3 15.0 19.6 21.3 21.6 33.7 7.0 22.0 31.6 117.3 148.0 13.0 18.7 17.0 21.8 20.2 21.8 24.9 7.5 22.0 32.4 143.4 87.0 11.7 20.0 19.0 15.0 21.0 19.8 21.0 8.0 24.0 34.3 161.6 43.0 16.9 23.0 20.0 15.0 21.0 16.7 19.0 8.5 28.0 29.1 131.8 25.0 17.2 25.0 15.0 15.0 21.0 22.5 17.0 9.0 32.2 31.0 147.0 22.8 17.4 27.0 15.0 18.0 18.0 19.4 17.0 9.5 26.4 40.0 121.0 21.7 16.7 34.0 12.0 23.0 20.0 17.2 28.0 10.0 27.5 40.0 52.0 21.5 18.2 39.2 13.0 20.0 20.4 23.0 28.0 10.5 27.7 33.0 86.0 23.3 19.3 67.4 14.0 16.0 19.5 17.0 31.0 11.0 30.8 28.0 113.0 23.2 16.5 104.5 13.0 15.0 20.6 19.0 33.0 11.5 25.2 27.7 105.7 22.5 17.6 27.7 14.0 15.0 21.7 30.0 30.0 12.0 23.4 27.6 127.6 22.4 18.9 17.9 12.0 19.0 19.9 30.0 29.0 12.5 24.5 29.5 144.5 22.6 18.3 16.0 13.0 15.9 21.2 26.8 23.0 13.0 21.7 36.3 109.3 21.2 18.6 16.0 13.0 16.7 19.4 17.6 24.0 13.5 23.9 40.2 79.2 20.1 19.0 17.0 13.0 18.5 19.5 20.5 26.0 14.0 28.0 77.6 95.0 19.8 20.3 18.0 13.0 16.3 20.7 24.3 24.0 14.5 27.0 114.3 148.8 19.7 24.6 17.0 13.0 15.2 19.9 22.1 23.0 15.0 27.0 71.3 159.6 20.5 25.2 17.0 15.0 16.1 21.0 25.8 23.8 15.5 27.0 48.2 164.4 24.4 24.3 17.0 15.0 20.3 20.0 20.7 22.6 16.0 27.0 37.1 125.1 20.2 22.5 19.0 15.0 21.4 22.0 18.5 28.4 16.5 28.2 46.0 150.0 21.0 23.6 23.0 16.0 16.6 21.0 22.4 26.2 17.0 29.4 51.0 129.0 22.0 25.8 24.0 14.0 17.8 20.0 25.2 17.5 27.6 49.0 130.0 23.0 24.2 21.9 14.0 17.0 20.1 22.0 18.0 24.7 47.0 185.0 23.0 22.3 22.7 13.0 24.0 20.3 26.0 18.5 27.9 39.0 202.0 22.0 21.5 25.5 13.0 26.0 18.5 29.0 19.0 28.6 48.8 270.8 21.0 24.7 22.3 11.0 22.0 24.7 21.0 19.5 27.5 53.7 342.7 22.0 25.9 20.1 14.0 16.0 20.8 23.0 Note: The sums of the column and row headings indicate the depth for the magnetic susceptibility readings. All magnetic susceptibility readings reported in 10-6 SI units.
108
TABLE A-4. PERCENT ORGANIC MATTER DATA FOR CORE 02
0 40 80 120 160 200 240 280 0 2.97 3.34 3.37 3.33 3.48 3.40 N.D. 1 2.36 3.27 3.01 3.53 3.25 4.02 3.74 N.D. 2 2.43 2.96 2.99 3.31 3.20 4.00 3.90 2.39 3 2.25 3.18 2.91 3.33 3.33 3.59 3.69 N.D. 4 2.25 3.25 2.93 3.65 3.43 3.45 3.79 N.D. 5 2.29 3.17 2.95 3.65 3.40 3.52 3.79 N.D. 6 2.40 3.24 2.74 3.39 3.35 3.69 3.92 N.D. 7 2.15 3.17 2.64 3.29 3.66 3.44 3.80 2.66 8 2.17 3.30 2.94 3.19 3.52 3.48 3.98 N.D. 9 2.17 3.15 2.95 3.20 3.41 3.80 3.92 3.63 10 2.49 3.40 3.31 3.40 3.17 3.65 4.18 3.93 11 2.74 3.08 3.07 3.88 3.28 4.08 4.45 3.43 12 2.84 2.84 3.04 3.45 3.03 3.87 3.76 3.59 13 2.95 2.99 2.96 3.50 3.46 3.62 3.71 3.56 14 3.96 2.95 2.93 3.35 3.43 3.43 3.50 3.64 15 3.36 3.03 2.95 3.39 3.19 3.71 3.01 3.53 16 2.62 3.25 2.95 3.33 3.41 3.79 3.42 3.41 17 2.89 3.15 2.89 3.51 3.16 3.62 3.96 3.50 18 2.78 3.36 2.91 3.27 3.17 3.51 3.92 3.49 19 2.69 3.26 3.18 3.44 3.48 3.43 4.54 3.55 20 2.93 3.49 3.15 3.36 3.39 2.96 4.08 3.59 21 2.56 3.00 2.81 3.38 3.57 3.88 3.82 3.54 22 2.82 3.23 2.97 3.46 3.64 3.61 3.70 3.80 23 2.66 3.41 2.96 3.30 3.32 3.58 3.50 3.44 24 3.00 3.93 3.11 3.45 3.05 3.61 3.53 3.89 25 3.45 2.88 3.12 3.68 3.08 3.47 2.86 N.D. 26 3.43 3.20 3.31 3.25 3.22 3.52 N.D. N.D. 27 3.04 3.19 3.30 3.31 3.51 3.06 N.D. N.D. 28 2.49 3.32 3.49 3.26 3.41 3.24 N.D. N.D. 29 2.92 3.07 3.33 3.44 3.34 3.30 N.D. N.D. 30 2.64 3.08 3.44 3.32 3.63 3.60 N.D. N.D. 31 2.61 3.14 3.21 N.D. 3.55 3.68 N.D. N.D. 32 2.73 3.35 3.53 3.36 3.44 3.43 N.D. N.D. 33 2.41 3.31 3.49 3.27 3.73 3.36 N.D. N.D. 34 3.22 3.32 3.77 3.32 3.46 3.70 N.D. N.D. 35 2.84 3.30 3.71 3.29 3.62 3.90 N.D. N.D. 36 3.26 3.00 3.71 3.46 3.38 3.86 N.D. N.D. 37 2.83 3.04 3.63 3.59 3.49 4.12 N.D. N.D. 38 3.07 3.11 3.34 3.84 3.82 3.65 N.D. N.D. 39 2.72 3.30 3.61 3.01 3.70 3.51 N.D. N.D.
109
300.0 340.5 380.5 420.5
0.0 N.D. 4.61 3.88 1.86 1.5 N.D. 4.51 3.81 2.07 2.5 N.D. 4.59 3.80 1.87 3.5 N.D. 4.63 3.89 1.55 4.5 N.D. 4.51 3.72 1.87 5.5 3.93 4.41 3.10 1.74 6.5 3.66 4.62 3.48 1.44 7.5 3.78 4.28 3.27 1.43 8.5 3.89 4.31 3.39 2.38 9.5 3.82 4.24 3.12 2.26 10.5 3.64 3.19 3.36 1.74 11.5 3.74 4.39 3.01 12.5 3.91 4.38 2.94 13.5 3.95 4.38 3.00 14.5 3.91 4.43 2.95 15.5 3.63 4.40 2.99 16.5 3.75 4.40 3.00 17.5 3.85 4.37 2.61 18.5 3.74 4.38 2.64 19.5 3.93 3.67 2.75 20.5 3.93 4.12 2.59 21.5 4.07 4.30 3.13 22.5 4.16 4.23 2.80 23.5 4.05 4.04 2.57 24.5 4.06 4.05 2.69 25.5 4.23 4.03 2.58 26.5 4.39 4.27 2.44 27.5 4.44 3.89 2.54 28.5 4.29 4.00 2.77 29.5 4.35 4.04 2.81 30.5 4.42 3.95 2.54 31.5 4.44 3.88 2.31 32.5 4.69 3.75 1.97 33.5 4.28 4.01 2.73 34.5 4.39 3.82 2.46 35.5 4.54 4.07 2.60 36.5 4.74 3.83 2.42 37.5 4.63 4.11 2.40 38.5 4.72 4.07 2.41 39.5 4.64 3.60 2.12 Note: The sums of the column and row headings indicate the depth for the loss-on-ignition measurements.
110
TABLE A-5. BIOGENIC SILICA DATA FOR CORE 02
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 0.60 1.75 0.23 0.80 2.00 0.82 1.00 1.92 1.35 1.20 2.47 2.05 1.40 2.06 1.99 1.60 2.68 2.94 1.80 2.67 3.30 2.00 2.92 4.01 2.10 3.23 4.39 2.40 2.44 3.14 2.60 2.47 3.04 2.80 2.66 3.12 3.00 2.35 2.62 3.20 2.32 2.59 3.40 2.10 2.32 3.60 2.37 2.61 3.80 1.99 2.17 4.00 2.72 2.93 4.20 2.37 2.61 4.40 2.30 2.58 4.60 2.14 2.45 4.80 1.90 2.20 5.00 2.23 2.61 5.20 2.08 2.42 5.40 1.99 2.29 5.60 2.04 2.32 5.80 2.22 2.49 6.00 2.03 2.24 6.20 1.62 1.63 6.40 2.35 2.21 6.60 2.36 2.06 6.80 2.01 1.62 7.00 2.42 1.80 7.20 1.64 1.23 7.40 2.37 1.74 7.60 1.78 1.28 7.80 2.23 1.56 8.00 2.07 1.41 8.20 1.97 1.34 8.40 1.81 1.24 8.60 2.01 1.37 8.80 1.90 1.29 9.00 1.86 1.26
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 9.20 1.75 1.16 9.40 2.01 1.30 9.60 2.06 1.31 9.80 2.00 1.25 10.00 2.20 1.34 10.20 2.09 1.18 10.40 1.99 1.04 10.60 1.81 0.87 10.80 1.70 1.77 11.00 1.64 1.63 11.30 1.76 1.69 11.60 2.08 1.92 11.80 2.05 1.84 12.00 2.03 1.78 12.40 1.96 1.76 12.60 2.62 2.38 12.80 1.86 1.71 13.00 2.03 1.89 13.20 2.20 2.18 13.40 2.09 2.19 13.60 1.93 2.14 13.80 2.00 2.34 14.00 1.88 2.32 14.20 1.96 2.36 14.40 1.96 2.29 15.75 1.88 1.79 16.25 2.07 1.97 16.75 2.12 2.26 17.25 1.85 1.97 17.75 2.14 2.11 18.25 2.11 2.09 18.75 2.25 2.58 19.25 1.63 1.87 19.75 2.20 2.64 20.25 2.09 2.50 20.75 1.93 2.11 21.25 1.78 1.94 21.75 1.81 1.94 22.25 1.88 2.01 22.75 2.06 2.08 23.25 1.47 1.49 23.75 2.17 2.24 24.25 1.84 1.90 24.75 2.47 2.39 25.25 2.19 2.11
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 25.75 1.70 1.87 26.25 1.57 1.73 26.75 1.70 1.27 27.25 1.49 1.12 27.75 1.89 1.75 28.25 2.09 1.95 28.75 2.32 2.23 29.25 2.20 2.12 29.75 1.91 1.49 30.25 2.04 1.59 30.75 1.81 1.64 31.25 2.07 1.88 31.75 2.02 2.12 32.25 1.85 1.95 32.75 1.96 1.37 33.25 1.79 1.25 33.75 2.08 1.92 34.25 1.97 1.82 34.75 1.89 1.88 35.25 1.58 1.58 35.75 2.00 1.74 36.25 1.98 1.72 36.75 1.73 1.78 37.25 2.99 3.07 37.75 1.91 1.98 38.25 1.54 1.60 38.75 1.82 2.11 39.25 1.92 2.23 39.75 1.92 1.91 40.25 1.90 1.89 40.75 1.73 1.48 41.25 1.82 1.56 41.75 1.80 1.63 42.25 1.69 1.53 42.75 1.88 1.54 43.25 1.73 1.42 43.75 1.86 2.43 44.25 1.81 2.36 44.75 1.89 1.47 45.25 1.89 1.46 45.75 1.86 1.59 46.25 1.95 1.67 46.75 1.63 1.66 47.25 2.25 2.29 47.75 1.93 2.05
112
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 48.25 1.95 2.06 48.75 1.82 2.13 49.25 1.96 2.30 49.75 1.47 2.06 50.25 1.71 2.39 50.75 1.91 1.79 51.25 1.88 1.76 51.75 1.78 1.72 52.25 1.76 1.71 52.75 2.04 2.18 53.25 1.61 1.72 53.75 1.69 2.32 54.25 1.64 2.25 54.75 1.74 1.67 55.25 1.85 1.77 55.75 1.53 1.79 56.25 1.56 1.83 56.75 1.64 1.86 57.25 1.28 1.45 57.75 2.25 2.16 58.25 1.76 1.70 58.75 1.91 2.26 59.25 1.59 1.88 59.75 1.57 1.83 60.25 1.81 2.11 60.75 1.98 2.04 61.25 1.88 1.93 61.75 1.79 1.72 62.25 1.91 1.84 62.75 1.88 2.35 63.25 1.89 2.37 63.75 1.79 2.16 64.25 1.36 1.64 64.75 1.85 2.01 65.25 1.74 2.25 65.75 1.71 1.15 66.25 1.76 1.19 66.75 1.74 1.34 67.25 1.37 1.05 67.75 1.76 1.96 68.25 2.06 2.29 68.75 2.09 2.24 69.25 1.78 1.91 69.75 1.96 2.05
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 70.25 1.95 2.04 70.75 2.02 2.36 71.25 1.80 2.10 71.75 1.34 1.41 72.25 1.30 1.36 72.75 2.31 1.65 73.25 1.40 1.00 73.75 2.07 1.85 74.25 2.11 1.89 74.75 0.95 0.88 75.25 1.82 1.68 75.75 1.56 1.65 76.25 2.02 2.14 76.75 1.97 1.88 77.25 1.84 1.75 77.75 1.82 1.42 78.25 2.15 1.69 78.75 1.79 1.90 79.25 1.97 2.09 79.75 1.90 2.55 80.25 1.87 2.51 80.75 2.17 2.24 81.25 1.84 1.90 81.75 2.05 2.87 82.25 1.76 2.47 82.75 1.81 2.07 83.25 1.90 2.17 83.75 1.96 2.04 84.25 1.86 1.94 84.75 1.84 2.14 85.25 1.73 2.01 85.75 1.98 1.72 86.25 2.08 1.81 86.75 2.09 2.22 87.25 2.20 2.35 87.75 1.96 2.37 88.25 1.94 2.35 88.75 1.85 2.42 89.25 1.95 2.54 89.75 2.00 1.73 90.25 2.18 1.88 90.75 1.97 1.82 91.25 2.06 1.90 91.75 2.01 1.59 92.25 2.20 1.74 92.75 2.12 2.54
Depth BSi BSi Flux s(cm) (%) (mg·cm-2
·yr-1) 93.25 2.09 2.51 93.75 2.16 2.84 94.25 2.01 2.65 94.75 2.26 3.05 95.25 1.69 2.29 95.75 2.65 3.00 96.25 2.76 3.13 96.75 2.46 3.09 97.25 2.41 3.02 97.75 2.44 2.87 98.25 2.81 3.30 98.75 2.76 2.74 99.25 2.54 2.52 99.75 2.23 2.22 100.25 2.46 2.45 100.75 2.64 2.63 101.25 2.52 2.51 101.75 2.66 3.10 102.25 2.39 2.78 102.75 2.51 2.85 103.25 2.49 2.82 103.75 2.40 2.83 104.25 2.15 2.53 104.75 2.37 2.34 105.25 2.70 2.66 105.75 2.95 2.96 106.25 2.89 2.90 106.75 2.71 2.92 107.25 2.51 2.70 107.75 2.75 2.36 108.25 1.57 1.34 108.75 2.14 2.41 109.25 2.54 2.86 109.75 2.55 2.30 110.25 2.65 2.38 110.75 2.51 2.68 111.25 2.12 2.25 111.75 2.63 2.86 112.25 2.68 2.91 112.75 2.60 2.47 113.25 2.66 2.53 113.75 2.71 2.31 114.25 2.59 2.21 114.75 2.16 2.54 115.25 2.11 2.47 115.75 1.74 2.48
111
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Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 116.25 1.90 1.79 117.25 1.56 1.76 117.75 1.60 1.58 118.25 1.82 1.79 118.75 1.64 1.80 119.25 1.84 2.03 119.75 1.88 2.01 120.25 1.81 1.94 120.75 1.54 1.64 121.25 1.40 1.50 121.75 1.50 1.44 122.25 1.46 1.40 122.75 1.55 1.49 123.25 1.98 1.90 123.75 1.65 1.33 124.25 1.52 1.23 124.75 1.72 1.58 125.25 2.60 2.39 125.75 1.90 1.83 126.25 1.83 1.77 126.75 1.81 1.41 127.25 1.65 1.29 127.75 1.79 2.00 128.25 1.80 2.00 128.75 1.75 1.59 129.25 1.69 1.54 129.75 1.68 1.79 130.25 1.75 1.87 130.75 1.73 1.46 131.25 1.67 1.41 131.75 1.77 1.65 132.25 1.85 1.72 132.75 1.90 1.74 133.25 1.83 1.67 133.75 1.75 1.10 134.25 1.70 1.07 134.75 1.67 1.55 135.25 1.68 1.56 135.75 1.03 0.95 136.25 1.09 0.97 136.75 1.04 0.82 137.25 1.41 1.10 137.75 1.22 1.42 138.75 1.29 1.53 139.25 1.17 1.38 139.75 1.51 1.29 140.25 1.17 1.00 140.75 1.25 1.41
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 141.25 1.45 1.63 141.75 1.42 1.21 142.25 1.32 1.12 142.75 1.13 1.03 143.25 1.24 1.12 143.75 1.08 0.82 144.25 1.19 0.90 144.75 1.21 1.26 145.25 1.33 1.38 145.75 1.33 1.80 146.25 1.29 1.75 146.75 1.15 0.95 147.25 1.26 1.04 147.75 1.27 1.29 148.25 1.23 1.25 148.75 1.42 1.26 149.25 1.32 1.17 149.75 1.47 1.19 150.25 1.45 1.16 150.75 1.35 1.09 156.25 1.29 1.03 161.25 1.67 1.21 166.25 1.33 1.17 171.25 1.76 1.32 176.25 1.84 1.42 181.25 1.72 1.41 186.00 1.92 1.39 191.25 1.76 1.53 196.25 1.76 1.19 201.25 1.53 1.28 206.25 1.91 1.52 211.25 2.19 1.45 216.25 2.30 1.72 221.25 1.59 1.12 226.25 2.20 1.22 231.25 2.18 2.17 236.25 3.06 2.28 241.25 4.09 2.30 246.25 4.07 1.99 251.25 4.86 3.60 256.25 3.39 1.82 260.25 3.80 2.71 260.75 5.26 3.83 261.25 5.41 3.94 261.75 4.33 4.03 262.25 3.87 3.60 262.75 4.68 2.86 263.25 4.57 2.79
Depth BSi BSi Flux (cm) (%) (mg·cm-2
·yr-1) 263.75 3.17 2.65 266.25 3.79 N.D. 288.25 2.75 0.42 289.25 2.89 0.44 290.25 1.85 0.25 291.25 1.71 0.37 296.25 1.62 0.29 301.25 1.23 0.25 309.75 1.77 0.20 314.75 1.36 0.23 319.75 1.85 0.32 324.75 1.87 0.39 329.75 3.84 0.69 330.75 6.00 0.93 334.75 2.08 0.38 339.75 1.92 0.31 341.75 2.04 0.36 344.75 2.21 0.39 345.75 2.55 0.37 354.75 2.36 0.38 359.75 2.08 0.45 364.75 2.26 0.46 369.75 2.05 0.40 374.75 2.04 0.63 377.75 1.74 0.22 379.75 1.93 0.55 382.75 2.01 0.42 384.75 1.67 0.37 389.75 1.48 0.33 394.75 1.72 0.37 397.75 1.24 0.27 399.75 0.99 0.28 404.75 1.15 0.21 407.75 1.75 0.48 409.75 1.18 0.21 414.75 1.27 0.27 419.75 0.72 0.11 424.75 0.58 0.14 429.75 0.90 0.28 434.75 0.91 0.28 436.75 0.65 0.27
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